NYSGA 2006
New York State Geological Association
78th Annual Meeting
Hosted by:
University at Buffalo Rock Fracture Group
Department of Geology, 876 NSC
University at Buffalo
Buffalo, NY 14260
R.Jacobi, General Chair and Compiler
Conference Headquarters:
Adam's Mark Buffalo-Niagara Hotel
120 Church Street
Buffalo, NY 14202
Dr. Alan I. Benimoff, Executive Secretary
NYS Geological Association
Department of Engineering Science and Physics,
The College of Staten Island/CUNY
2800 Victory Boulevard
Staten Island, NY 10314
benimoff@postbox.csi.cuny.edu

A1: FAULTING AND MINERALIZATION IN THE CAMBRO-ORDOVICIAN SECTION OF THE MOHAWK VALLEY
Paul Agle, R.Jacobi, Charles Mitchell
UB Rock Fracture Group
Department of Geology
University at Buffalo
Buffalo, NY 14260
Rich Nyahay, Brian Slater, Langhorne Smith
Reservoir Characterization Group
New York State Museum
3140 Cultural Education Center
Albany, NY 12230
Sponsored in part by NYSERDA (New York State Energy Research and Development Authority)
We thank STATEMAP and NYSERDA for funding research reported herein.
INTRODUCTION
The Cambro-Ordovician section in the Mohawk Valley region is one of the most well known and longest studied Cambro-Ordovician sections in the world. From Amos Eaton in the 1820s (e.g., 1824) to Vanuxem (1834, 1838, 1843) to Emmons (1842), Schuchert (e.g., Clarke and Schuchert, 1899) to Ruedemann (1897, 1925) to Cushing (1905) to Kay (1937, 1968) to the more recent geologists (e.g., Bird and Dewey, 1970; Fisher, 1977, 1979; Fisher et al., 1970; Landing et al., 1996; Jacobi 1981, 2002; Bradley and Kidd 1991; Bradley and Kusky, 1986; Cisne and Rabe, 1978; Cisne et al., 1982; Mitchell et al., 1994; Goldman et al., 1994; Lehmann et al., 1995; Brett and Baird 2002; Baird and Brett 2002), and many, many more (see review in Jacobi and Mitchell, 2002), the Cambro-Ordovician geology in the Mohawk Valley (Fig. 1) has provided critical data for understanding Cambro-Ordovician stratigraphy (Fig. 2) and the inferences that can be drawn from the stratigraphy and structure. For example, the original application of plate tectonics to land geology utilized the Mohawk Valley geology (Bird and Dewey, 1970; Fig. 3A), as have revisions to the original model (e.g., Jacobi, 1981; Fig. 3B; Bradley and Kidd, 1991; Fig. 3C). More recently, the Mohawk Valley section has functioned as the test case for detailed volcanic ash geochronology (tephrachronology) and graptolite correlations (e.g., Mitchell et al., 1994), and for gauging the relative importance of tectonic versus eustatic sea level signals on the apparent sequence stratigraphy in the basin (e.g., Mitchell et al., 1994; Lehmann, 1994; Joy et al., 2000; Jacobi and Mitchell, 2002; Brett and Baird, 2002; Baird and Brett, 2002).
During the past decade, the Mohawk Valley section has been used as an analogue for subsurface hydrocarbon reservoirs in the Cambro-Ordovician section. The major gas discoveries in the Trenton/Black River in central NYS (and elsewhere) have driven a renewed interest in the exposed Mohawk Valley Cambro-Ordovician section. The present attention involves determining and understanding the factors that controlled the development of the Cambro-Ordovician gas reservoirs.
The present field trip examines a spectacular quarry floor that is a probable analog for the dolomitized Trenton/Black River reservoirs in central NYS. Trenches, cores, GPR, and water blasting of the outcrop led to an integrated, very comprehensive picture of how secondary porosity and dolomitization develop. The field trip also visits faults that may have been the pathways for the fluids that dissolved the limestone, resulting in vuggy porosity, and that precipitated hydrothermal dolomite and other minerals in the void space. The trip includes several optional outcrops that display the stratigraphy and the characteristic porosities in various units. Finally, the trip stops at the well-known NYS Thruway locality where the Thruway Disconformity is exposed, separating ribbon limestones of the Dolgeville Formation from the overlying black shale of the Utica (Indian Castle Member). This last stop is particularly interesting for those who have an interest in the black shales of the Utica. Because of severe time constrains, the trip cannot stop at all the localities, but we have included the optional stops for geologists to visit at a later date. Here we can examine in outcrop the stratigraphy, faults, fluid migration along these faults, and tectonics in the Mohawk Valley.
This field trip is the outgrowth of the work of two research groups, the NYS Museum Reservoir Characterization Group (e.g., Smith and Nyahay, 2005), and the UB Rock Fracture Group (e.g., Jacobi and Mitchell, 2002). Each group has worked for some time in the Mohawk Valley; the NYS Museum Reservoir Characterization Group has investigated a hydrothermal dolomite reservoir analog in the Palatine Bridge Quarry (STOP 1); their integrated technologies yielded a significant understanding of the hydrothermal systems. The Reservoir Characterization Group led a workshop and field trip to this site, (Smith and Nyahay, 2005). The UB Rock Fracture Group and ITAS (Integrated Tectonics and Stratigraphy Group at UB) have been working on stratigraphy and structure farther west in the western part of the Mohawk Valley sequence for 25 years (albeit sporadically; e.g., Ritter, 1983; Mitchell et al., 1994; Jacobi and Mitchell, 2002).
Below we first review the Mohawk Valley Cambro-Ordovician stratigraphy that incorporates the units and structures we will inspect on this field trip. Overall, the Cambro-Ordovician sequence in the Mohawk Valley records a Laurentian passive margin carbonate bank that is overstepped by black shales and later coarser clastics. The overstep resulted primarily from a continent/island arc collision wherein the Laurentian continental passive margin first passed over the peripheral bulge and then dropped down into an east-dipping subduction zone (e.g., Jacobi, 1981; Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985). A eustatic sea level rise may have been superimposed upon the tectonic signal (e.g., Brett and Baird, 2002). We also review the structure of the region, and the concepts involved in the dolomitization and development of the porosity typical of the Cambro-Ordovician carbonate section. A key component of the trip will be the link between faulting, fluid flow and reservoir development.

A1: FAULTING AND MINERALIZATION IN THE CAMBRO-ORDOVICIAN SECTION OF THE MOHAWK VALLEY
P.Agle, R.Jacobi, C.Mitchell, R.Nyahay, B. Slater, L.Smith

STOP 1. Palantine Bridge Quarry
The Palatine Bridge study site is located in an inactive quarry along the Mohawk Valley, approximately 50 miles west of Albany, NY. The outcrop occurs as a highly localized dolomite alteration in the Tribes Hill limestone that makes up the floor of the quarry. Although there are at least three areas of the quarry that have been dolomitized, research has been focused on the one region most accessible. The site consists of two dolomite bodies (Figs 7 &8). The eastern most body (body 1) is a long, linear feature approximately 55 feet long, 5 feet wide, and has a strike of 305°. The second body (body 2) is located immediately west of the first, and is also a linear feature with similar strike, however this body is substantially longer than the first, measuring 110 ft. Body 2 is broken into two parts which are separated by a southerly bend, referred to as the jog. The dolomite is easily distinguished from the surrounding limestone by its distinct orange color and massive texture.
Analog Comparison
Although it is much smaller than the hydrothermal dolomite fields in production today, all other characteristics of the quarry are virtually identical to those of its larger counterparts (Fig. 9). Hydrothermal dolomite fields such as the Rochester field of Ontario, CAN, the Albion-Scipio fields of Michigan, and the numerous gas fields of South Central NY all bear striking resemblance to the Palatine Bridge Quarry outcrop.
Setting
The Palatine Bridge Quarry lies within the Tribes Hill unit of the Upper Beekmantown Group. It is separated from the Precambrian basement by the Little Falls dolostone (lower Beekmantown Group) and the Potsdam sandstone which may or may not be present as it thins to the east. The Tribes Hill is a thin to medium-bedded limestone which was deposited in a shallow, peritidal to subtidal environment. It is underlain by the Late Cambrian Little Falls dolostone which is described as a thick-bedded dolostone and is also believed to have formed by deposition in a shallow, peritidal to subtidal environment. The Little Falls is approximately 400ft. thick in this portion of the Mohawk Valley.
The Tribes Hill Limestone outcrops in Palatine Bridge; however, further south it is unconformably overlain by the Upper Ordovician Trenton- Black River Group and Utica Shale. The Trenton and Black River units are well known for their oil and gas plays along the eastern portion of the US and into southern Ontario. The majority of these plays are contained in hydrothermal dolomite reservoirs.
The Beekmantown Group was deposited during the Late Cambrian and Early Ordovician along a passive margin following the opening of the Iapetus Ocean. However, deposition of the Middle Ordovician Trenton- Black River Group has been demonstrated to have been fault controlled. Active tectonism during the Middle-Late Ordovician is associated with the Taconic Orogeny in which Proto-North America collided with the Taconic Island Arc terrain. Flexure of the North American plate during its attempt to be subducted under the island arc led to extensional faulting and produced a series of north east-south west striking normal faults, some with offsets of over 1000 ft.
Field Relations
There are many characteristics of the outcrop that indicate a fault related hydrothermal origin. The central sag flanked by monoclines is a typical feature of many hydrothermal structures (Fig. 10). Fault breccia and vuggy porosity can be observed in several places (Fig. 11). These areas are commonly located near, but not limited to, the tips of each body. An intense fault and fracture zone surrounds the entire outcrop increasing the width of the affected area to approximately 15 ft. (Fig. 12). Many faults have vertical slickenlines which indicate a stress regime that was predominantly extensional. Groups of faults commonly join to form a relay ramp structure as seen in Figure 13.
Ground Penetrating Radar
In the spring of2004 a shallow ground penetrating radar survey was run over body 2. The purpose of this survey was to aid in the excavation process and to get a profile view of the outcrop. Although the radar was only able to penetrate a couple meters, it showed that the outcrop has geometry very similar to that of the Trenton / Black River gas plays (Fig. 14). Each body consists of a central syncline, or sag, that is flanked by smaller anticlines on either side.
Cores
A series of six cores 40 - 60 ft. deep were taken from strategic points on the outcrop so that cross sections could be constructed both along strike and perpendicular to the bodies. The perpendicular cross section (Fig. 15) demonstrates that the dolomitization is highly localized as Hole Six, less than 20 ft. outside the body, is almost completely undolomitized. The parallel cross section (Fig. 16) shows that the limestone gap between bodies 1 and 2 is only 3 feet thick. The two bodies join at depth.
Trenches
Six trenches measuring 3' wide by 15'long by 18" deep have been cut across the outcrop (Fig. 17) to expose the folds of the body as well as the dip and offset of the fractures (Fig. 18).
Paragenic Sequence
Examination of thin sections made from outcrop and core samples indicate a fairly complicated diagenetic history including multiple episodes of faulting and mineral precipitation (Fig. 19).
Geochemistry
Strontium isotope analyses done on dolomite samples yield values significantly higher than that of seawater during the time of Tribes Hill deposition (Fig. 20 & 21 ). Limestone and calcite samples plot on the curve as expected. The elevated strontium content is interpreted to be caused by interaction of the parent fluid with siliciclastic rocks, such as the Grenville basement, for an extended period of time.
Stable Isotope analyses show that dolostone crystals formed from a fluid much lighter than seawater dolomites (Fig. 22). It is believed that these lighter isotope ratios are an indication of higher temperature fluids traveling upward from deeper, hotter depths.
Fluid inclusion analyses conducted on 40 samples indicate that both dolomite and calcite formed at temperatures greater than 100°C (Fig. 23). Although thermal chronology studies indicate that the Tribes Hill has been buried to a geothermal temperature of over 100°C, core analysis and paragenic sequencing suggest that dolomitization occurred at a shallow depth before significant burial.
P.Agle, R.Jacobi, C.Mitchell, R.Nyahay, B. Slater, L.Smith

STOP 2. Ingham Mills
Ingham Mills is a well known paleontological and stratigraphic locality due to the complete Trenton/Black River section at this location (Fig. 24). The section here is over 23 m thick and is the farthest eastward extent of the Watertown and Napanee Formations. The section is exposed on the spillway beneath the hydro dam that holds back the waters of Keyser Lake. The lowermost outcrop at the base of the spillway is a buff-weathering, fine-grained dolostone. This dolostone has previously been classified as the uppermost surface of the Little Falls Formation (Cameron et al., 1972; Cameron and Kamal, 1977; Mitchell and Marsh, unpublished data 1998). The hummocky nature of the surface does resemble thrombolitic horizons within the Little Falls Formation, though the fine-grained texture of the dolostone is very rare within this formation. However, new observations reveal that this is not a simple sedimentary contact. The dolomite observed at the section is actually an aureole of dolomitization hosted within the Lowville limestone along a fracture trending 290° (Fig. 25). On the north side of this fracture, the limestone layers clearly "onlap" over the dolomite body, but on the south side of the fracture, the dolostone grades rather abruptly into limestone within the same bedding plane. This gradation indicates that this dolostone was formed by migration of a dolomitizing fluid through the limestones along this fracture trend. The geometry of this dolomitization aureole and the surrounding limestones are strikingly similar to those observed at the Palatine Bridge Quarry (STOP 1).
Above the dolomite body another 8.66 m of Lowville Formation is exposed. The rocks here consist of blue-gray to dove-gray, very fine-grained limestone. Many horizons are riddled with vertical burrows of Phytopsis tubulosum and vertical fingers of anthraxolite (filling in voids of dissolved burrows?). One layer in the upper Lowville contains large colonies of Tetradium both over-turned and in life position. Two horizons within the Lowville show distinct tidal channel structures. The lowermost of these horizons show a rubbly lag on the lower surface of the channel. The rocks herein assigned to the Lowville Formation have been previously designated as Gull River by Cameron and others (1972), whereas Cornell (2003) classified the lowermost 3 m of this section as the Pamelia Formation and assigned the remainder to the Lowville Formation. At the top of the Lowville Formation lies 0.35 m of dark gray, rubbly limestone which we assign to the Watertown Formation. These rocks have also been called the Selby Formation by Cornell (2003). Above this horizon is 3.42 m of Napanee Formation (the lowermost formation of the Trenton Group). These rocks are fine-grained dark gray limestones that weather to a distinctive chocolate brown color. The limestones also contain abundant dark gray, calcareous shale interbeds. A prominent deformation feature of enigmatic origin is located on the west side of the outcrop; this feature incorporates an unknown amount of the upper Lowville, the entire thickness of the Watertown Formation, and the lowermost 2 m of the Napanee Formation. This structure is probably the result of karst/dissolution collapse in the Lowville Formation. The overlying 9.3 m of the Kings Falls Formation is differentiated from the Napanee Formation by its coarser-grained texture and large ripples indicating a higher energy depositional environment. A few of the rippled horizons contain large clasts and one bed at 14.2 m contains clasts of some underlying units including gneiss pebbles derived from the Precambrian basement. These clasts had to have been shed from a nearby fault scarp that must have existed nearby during the deposition of these limestones. The inferred position of the Dolgeville Fault lays a mere 400 m to the west of Ingham Mills and may well have been active at this time (Jacobi and Mitchell, 2002). The uppermost 1.2 m of section belong to the Sugar River Formation, which is differentiated from the underlying Kings Falls by a decreased grain size, absence of high-energy ripples, and numerous bedding planes with abundant Prasopora colonies.
P.Agle, R.Jacobi, C.Mitchell, R.Nyahay, B. Slater, L.Smith

STOP 3.
The power dam at Dolgeville on East Canada Creek is the most accessible exposure of the Dolgeville Fault (Fig. 26). The fault plane itself is not visible at this outcrop, though the effects of faulting are clearly and dramatically observed in the downthrown block. The fault disappears beneath thick glacial cover to the north and cannot be traced reliably for any great distance. To the south, the actual fault plane outcrops along the east bank of Kyser Lake (now accessible only by boat) and consists of a well-cemented fault breccia with iron/sulfide mineralization (Fig. 27). Cushing (1905) drew an excellent cross-section of this southern outcrop along East Canada Creek as it was before the dam at Ingham Mills raised the river level to create Kyser Lake (Fig. 28). The fault does not outcrop south of this location, but stratigraphic evidence reveals that it probably crosses the southern part of the lake and continues southward (with diminishing throw), where it passes somewhere between West Crum and East Canada Creek. If the fault crosses the Mohawk River, then the throw is severely reduced. At the Dolgeville Dam, the fault trends roughly N-S and drops down to the west (antithetic to the other major faults). Cushing (1905) estimated the throw of the fault at 300 ft (91.4 m), whereas Fisher (1954) estimated 450 ft (137.2 m). The Dolgeville Fault here exhibits dip slip kinematic indicators. Minor drag folds have nearly horizontal axes, and down-dip slickensides (Fig. 26 inset) occur on minor faults and on bedding around some of the folds (Jacobi and Mitchell, 2002). Farther south Bradley and Kidd (1991) also found down dip slickensides on a minor fault near the Dolgeville Fault (their Figure 8a). Thus, if these kinematic indicators are related to Ordovician motion, then the Dolgeville Fault was primarily a dip slip (normal) fault. We believe this fault originated in response to plate stress as it descended into the trench. The fact that this fault cross-cuts gravity anomalies and that the N-trending fractures (Dolgeville trend) usually post-date the NNE trending fractures (Little Falls trend) indicate that slip on this fault likely initiated relatively late as a result of peripheral bulge tectonics and is younger than the Little Falls Fault. There are a few folds and faults that indicate a later reversal in motion on the Dolgeville fault.
At this site, the up thrown block consists of sub-horizontal Little Falls dolostone that outcrops along Dolge Ave. overlooking the falls. Looking northeast at the spillway outcrop, the down thrown block consists entirely of Indian Castle black shale with rare limestone beds. The beds of the falls are sub-horizontal near the powerhouse, but slowly begin to dip to the west nearing the fault. At the small line of trees, the bedding dip abruptly increases across a minor fault. The beds then continue to gradually increase in dip to angles measured as high as 59°W, although the beds are much steeper on the inaccessible shear cliff-face. Within the steeply dipping beds, the Dolgeville/Indian Castle contact is gradational, with an upsection increase in the thickness of the intervening shales. No time gap is thought to exist at this locality between the Dolgeville and the Indian Castle (Jacobi and Mitchell, 2002). The lack of a significant, observable "Thruway disconformity" (viewed at STOP 5) is consistent with the lack of a significant number of the distinct folds that characterize the NYS Thruway outcrop. No unequivocal slump folds, no time gap, and no locally angular unconformity all suggest that the "Thruway disconformity" is not a regional unconformity rather it is a slide scar, and the sediment slide did not affect this section (Jacobi and Mitchell, 2002). That scenario is consistent with the tectonic position of this outcrop - it is in the base of the Dolgeville graben, whereas the NYS Thruway outcrop is located on the horsts of the Little Falls fault system.

STOP 4: Outcrop at Intersection of Rts. 5 and 169
Park along wide shoulder near base of incline across from intersection with NY 169
Cushing (1905) described several outcrops of the Little Falls Fault in the field area. Today, these outcrops no longer exist or are severely degraded. The only locality that exposes the fault plane is the Thruway cut south of the Mohawk River (STOP 5). To the north of this locality, a strong NNE-trending topographic lineament is easily recognized. In the creek to the east of Moreland Park, a NNE-trending fracture intensification domain (FID) occurs within the Little Falls Formation along with loose blocks of stock-work breccia. The fractures within this FID have an average strike of 038° with a frequency of 20 fractures/m. This lineament appears to offset both the Precambrian-Little Falls contact and the Little Falls-Tribes Hill contact with the western side down-dropped. To the south, this topographic lineament truncates Moss Island on the west side. Across the Mohawk River, this lineament is clearly visible as a small cleft in the "Rollaway". The Precambrian-Little Falls contact appears to be offset down to the east across this cleft, though no other evidence for faulting was observed at that locality. The fact that the westerly block is down-dropped on the north side of the Mohawk River, whereas the easterly block is down-dropped on the south side indicates that this splay is either a scissors fault or there is an unrecognized cross fault between the two localities. At the present location, this lineament is expressed as the small creek that we see at the stop. At this site, we observe another NNE trending (average strike= 040°) FID (fracture frequency = 13.5 fractures/m) and also a fairly dramatic change in lithology in the Precambrian basement. The east side of the creek is well-foliated charnockitic gneiss, whereas the west side is darker, much more mafic, poorly-foliated gneiss. Slickenlines can be observed on both NNE and WNW surfaces at this outcrop (Fig. 29). Fractures of the NNE-trend increase in frequency nearing the fault (Fig. 30). The preferred growth orientations of crystals within the slickenlines indicate that the NNE trending fault experienced right lateral oblique slip with the SE fault block down-dropping (Fig. 29A). This sense of motion is opposite to the stratigraphic offset along this splay of the Little Falls Fault on the north side of the Mohawk River, thus the offset on the Little Falls Fault may be the result of several cumulative slip events with fault motion reversing at intervals throughout its history.
The slickenlines on the WNW surface are sub-horizontal (plunge= 04°) and indicate left-lateral slip with the northern block being slightly down-dropped (Fig. 29B). Leaving the creek and looking at the road cut exposures along Rt. 5, another WNW-trending surface shows well developed slickenlines. The slickenlines at this location record left-lateral oblique slip with a slight down on the south component (Fig. 29C). This sense of offset does not agree with the offset along the inferred WNW transfer fault to the north, but it is consistent with the sense of motion required to drop the Precambrian-Little Falls contact down on the south side of the Mohawk River (on an inferred WNW fault). On the north side of the river, this contact lies at 560 ft (170.7 m), whereas on the south it is at 400 ft (121.9 m). This offset can be partially explained by regional dip on this contact, which is estimated at 03° to the south based on the topographic gradient observed on the north side of the Mohawk River. Projecting this dip across the Mohawk River would account for only 125 ft (38.1 m) of offset. The kinematics of all the slickenlines (Fig. 29D) at this locality are consistent with the expected motions in an eastward directed far-field compressional regime, though these faults probably experienced several stages of slip and reversals in sense of motion.

STOP 5. Thruway cut South of Little Falls
Park in construction zone at northwest corner of Paradise Rd. and Thruway intersection
The NYS Thruway roadcut south of Little Falls provides a 1.2 mile long look at the ribbon limestones of the Dolgeville Formation, the overlying black shale of the Indian Castle Member, and the unique contact between these two units (Fig. 31 ). At the Thruway road cut, the upper few meters of the Dolgeville Formation are intensely folded. Fisher (1979) first studied these folds in 1953 when the roadcut was dug. Fisher (1979) noted the dominant westward vergence of the folds, and suggested that westward slumping away from the foredeep (not toward the basin) had caused these folds. He proposed that westward gravity sliding of the Taconic Giddings Brook "Slice" into the foredeep (Magog, or Snake Hill, Basin), which was located east of the thruway roadcut along an axis between the Hudson River/Lake Champlain and the Green Mountains/Taconics) had depressed the crust below the gravity slides and caused the Adirondack Arch to rise. Since the arch was located east of the roadcut, the rising arch would have reversed the original eastward slope to a westward slope, and thus the slumps traveled west, not east, off the carbonate ramp. The overlying Indian Castle Member lies largely undisturbed above the local unconformity known as the Thruway Discontinuity (after Baird and others, 1992).
Fisher's (1979) fold vergence was supported in general by studies by two of Jacobi's students who analyzed the outcrop in the 1980s (Ritter, 1983; Koslosky, 1986). Jacobi and Mitchell (2002) further analyzed the folds and paleoflow indicators at this site. Most of the folds are disposed in "trains" and affect only a relatively few beds; i.e., they are intrastratal (or intraformational), disharmonic folds. As described in Jacobi and Mitchell (2002):
Only kink band folds (box folds) affect more than a few beds. At the unconformity, eroded crestlines and troughs filled with breccia from the ribbon limestones, demonstrate that at least here the folds were exposed at the seafloor to erosive processes. These folds are not the result, therefore, of relatively deeply buried bedding-parallel slip. Both Fisher (1979) and Koslosky (1986) suggested that the folds were "slump folds" related to downslope slump processes, based on fold form and character. Highly contorted crestlines of some folds indicate that these folds underwent variations in strain along their axes as the sediments were transported downslope during the sliding event(s).
At this stop (near milepost 213), the asymmetric folds have fairly straight crestlines and west vergence. Application of the arc separation method (Fig. 32) yielded a down-paleoslope direction of 234.5° at the time of folding (Fig. 33). At a western outcrop (at milepost 214.8), the asymmetric folds exhibit a southwest vergence and the arc separation method yielded a down-paleoslope direction of 229° at the time of folding. These values are in accord with the detailed analyses of 698 folds by Koslosky (1986). Other features (Fig. 34) in the eastern outcrop that suggest a southwesterly-directed paleoslope at the time of sliding include the following observations:
1. small west-directed thrusts that affect single limestone layers in the Dolgeville Formation,
2. westward-directed intrastratal micro-fold nappe in the Dolgeville Formation,
3. imbricate, east-dipping clasts in breccias in the troughs of folds at the unconformity that suggest a westerly current flow immediately after the folding (sliding event[s]), and
4. a small slide scar that faces west.
Paleo flow data from the Dolgeville Formation at this outcrop imply a flow toward 207°; this direction is only 27° away from the paleoslope determined from the asymmetric folds. Graptolites and parting lineations above the unconformity show that the local paleoslope reversed directions after the time of slumping.
The major splay of the NNE-striking Little Falls Fault ("C") cuts the units in this stop (Fig. 35). The fact that bedding dip in the down thrown Utica (on the east side of the fault) is not perpendicular to the fault plane is evidence that this fault probably experienced oblique slip at this location (Jacobi and Mitchell, 2002). The newly recognized oblique-slip slickenlines along faults in the Precambrian basement at Little Falls (STOP 4 along Little Falls Fault "B") are consistent with the sense of rotation on the block inferred from the anomalous dip. In the past, no kinematic indicators were observed in the fault surface at the Thruway, but with the recent NYS Thruway slope abatement project, we may find kinematic indicators. Minor splays of the Little Falls Fault occupy valleys to the east and are recognized by minor changes in dip across the valleys. It is our contention that motion on these faults reversed the local paleoslope, causing the slumping activity recorded as the slump folds and the Thruway disconformity (a slide scar). This motion was related to the plate descending into the trench. (It should be noted that others maintain that the paleoslope did not reverse; Brett and Baird, 2002).

STOP 6. (Optional) Manheim Power Station
At the Manheim Power Station, the Manheim Fault is exposed along both sides of East Canada Creek (Fig. 36). Access from the Power Station driveway is on the west bank and permission is needed as this is private property. From this side of the creek, the rocks are easily accessed even at high water. The rocks along this bank superficially resemble the interbedded ribbon limestones and shales of the Dolgeville Formation and have been designated as such by previous workers (Megathlin, 1938; Fisher, 1954; Bradley and Kidd, 1991). However, faunal evidence indicates that the rocks observed here are the "Allen Road beds" of the Flat Creek Member (Mitchell, 2006 personal communication). Numerous fractures and veins dissect the outcrop at this locality. The fractures are of NE, WNW, and NNW orientations. It was expected that the NE-trending (fault-parallel) fractures would increase in frequency nearing the fault plane, but actually the NNW (fault-perpendicular) fractures show the most marked increase (Cross, 2004; Cross et al., 2004; Fig. 37). At the northern extent of the outcrop on this side of the creek is a fault breccia. This breccia contains lithologies similar to those of the adjacent Flat Creek Member consisting of large, rounded clasts of limestone that are well-cemented in a dark, calcareous mud matrix. Across the creek on the eastern side, one can see beds of the Flat Creek Member drag-folded upward on the eastern block to angles approaching 60°SE juxtaposed against gently dipping (07°SE) beds of the Little Falls Formation on the western block (Bradley and Kidd, 1991). Here the fault strikes roughly 50° (it is drawn as roughly N-trending in Fisher, 1980) and drops down to the east (Cross, 2004). According to early workers (Conrad, 1837; Vanuxem, 1838) there is a one inch thick calcite vein with associations of galena and pyrite present, which were mined for a brief period in the area of the fault. In addition, the fault zone hosts a peridotite dike. Similar dikes in the area roughly parallel the fault trend and have been dated using Rb-Sr as 130±10 million years old (Fisher, 1980). Estimates of the offset on this fault vary greatly with Darton (1895) claiming 152 ft (46.3 m), Megathlin (1938) 60 ft (18.3 m), Fisher (1954) 400 ft (121.9 m), and Bradley and Kidd (1991) 131 ft (39.9 m).

STOP 7. (Optional) Miner Rd. Outcrop
In the course of widening Miner Rd., an excellent outcrop of the lowermost Tribes Hill Formation was exposed. At this locality, the rocks are a light gray, medium-bedded (11 cm average), coarse-grained dolostone. The bedding at this locality trends 002/03W. Several beds show excellent herring-bone cross-stratification typical of a tidal environment. Four distinct fracture sets are present at this outcrop (N, NE, WNW, and NW). At this outcrop, the NE trending fracture set (average strike = 53°) is master (oldest) and parallel the inferred orientation of the Little Falls Fault just to the east. Reddish staining along the NW set is interpreted as sulfide and/or iron-rich vein mineralization (Fig. 38). The outcrop also hosts many vugs within the dolostone matrix. These vugs are variably filled with white calcite (Fig. 39) and occasional buff-colored dolomite. The presence of the veining and vug development are indicative that this horizon was favorable to the migration of fluids from the nearby fault. Heading south along Miner Rd. from this outcrop, there are excellent exposures of the upper Little Falls Formation in the woods on the western side of the road. These exposures contain stromatolite colonies and also host vugs partially filled with white calcite.

STOP 8. (Optional) Route 167 Overlook
The scenic overlook of Rt. 167 just before entering the city of Little Falls, NY provides an excellent vantage point to appreciate the magnitude of the fault offsets in the region (Fig. 40). Though the Little Falls Fault is not directly exposed at this locality, the steep linear scarp must be a close approximation of the fault plane. If one parks in the designated area and walks to the fence, a grand overview of the Mohawk River Valley opens up before you. In contrast to the narrows at Little Falls, NY, in the resistant Precambrian basement, the valley opens broadly to the east in the much less resistant shales of the Utica Group. If the gate is open, it is possible to walk out onto a small outcrop of Precambrian basement at the edge of the scarp. The foliation here is well defined and trends 284/26NE. This foliation is fairly uniform throughout the Little Falls area. Fractures in this outcrop trend NNE and WNW and are mutually abutting. The NNE striking fractures (average strike = 037°) of this outcrop parallel the topographic lineament and the inferred trend of the major splay ("C") of the Little Falls Fault in this area. While standing on this outcrop of Grenvillian basement, one can look down into the valley approximately 150 feet below and see the Thruway Interchange and its impressive exposures of the Upper Ordovician Indian Castle Member. Cushing (1905) estimates the throw on this fault at this location at 650-750 ft (198.1-228.6 m), whereas Fisher (1954) estimates 900 ft (274.3 m). The fault extends northward into the crystalline rocks of the Adirondack Dome and southward where it dies out in the Silurian shales. Across Rt. 167 from the scenic overlook is an old quarry. This quarry contains a 16 m section of the lowermost Little Falls Formation, though the lower contact with the Precambrian basement is not exposed here. This section of Little Falls dolostone is very coarse-grained and contains large amounts of quartz and some large detrital feldspar grains. Many beds display cross-bedding and a few isolated horizons host vuggy porosity.

STOP 9. (Optional) Little Falls Bike Path Area
The bike path on the south side of the Mohawk River now occupies an old railroad bed. Walking eastward from the designated parking area, large exposures of Precambrian basement can be observed. Approximately halfway through the cut about 20 ft from ground level on the south side, the contact between the Precambrian basement and the Little Falls Formation is exposed (Fig. 41). The contact is inaccessible without climbing gear, but clearly visible. Walking back to the parking lot and heading south, there is a footpath that winds its way up above the bike path. Along this path, there are excellent exposures of Little Falls dolostone. These outcrops are buff-weathering, thickly-bedded, coarse-grained dolostone that display crossbedding, stromatolites, and numerous vugs. These vugs are filled with saddle dolomite, quartz, and anthraxolite (Fig. 42). Dolomite usually lines the vugs, followed by anthraxolite, though a few vugs contain anthraxolite films that pre-date the dolomite. Where quartz is present, it is always the last mineral to precipitate, although anthraxolite commonly occurs as solid inclusions within the crystals. The fact that the age relationships of dolomite and anthraxolite are inconsistent, suggests that these minerals may have significant overlaps in their time of precipitation. Dunn and Fisher (1954) analyzed the anthraxolite from three similar vugs in the area and obtained an average composition of 90.39% carbon, 7.03% water, and minor components of nitrogen, sulfur, and ash. The temperature of precipitation of these vug-filling minerals is poorly studied. Dunn and Fisher (1954) cite Dr. F.G. Smith as having obtained a temperature of 51 °±2°C for a quartz crystal from the formation. O'Reilly and Parnell (1991), based on limited analysis of bitumen reflectance in the anthraxolite and fluid inclusions in calcite and quartz, developed a four-stage model in which the first stage (70°C) deposited calcite cements, the second stage precipitated calcite veins and vug fills (70 to 130°C), a third stage of anthraxolite (and presumably dolomite) and quartz (120°C), and a fourth very high temperature stage 200±10°C. Further fluid inclusion analyses of vein fill in the study area is currently underway.

STOP 10. (Optional) Nowadaga Creek
The exposures along Nowadaga Creek at the River Rd. bridge offer an opportunity to study the Dolgeville Formation and the contact with the overlying Indian Castle Member. The creek bed downstream of the bridge exposes the upper surface of the Dolgeville Formation and is readily observable at low water levels. The upper beds of the Dolgeville Formation are a complexly interference-folded surface (Fig. 43) and the contact with the overlying Indian Castle Member is very abrupt. For a detailed study of the folds and paleoflow directions at this locality see Jacobi and Mitchell (2002). Downstream from the contact is exposed a nearly complete section of the Dolgeville Formation. Very small (1O cm wavelength) asymmetric folds with irregularly-bearing axes are the result of mini-blind thrusts involving single limestone beds (Marsh, 2000). These thrusts suggest that the entire Dolgeville section was deposited during times of instability. The fractures present within the Dolgeville trend NNE, WNW, and NW. There are also veins present in the Dolgeville Formation at this locality. These veins parallel the NNE fracture trend and range in thickness from 1-3 mm of white calcite. Upstream of the bridge, the creek is floored by the Indian Castle Formation. This abrupt transition from contorted Dolgeville ribbon limestones and shales to undisturbed calcareous black shales represents a significant time gap, but not as much as that at the Thruway Discontinuity (Jacobi and Mitchell, 2002). In an outcrop on the northern bank, the lower Indian Castle Member is seen to contain a few thin carbonate beds, but overall the lithology is dominated by black calcareous shale. The fractures hosted within these shales show three prominent orientations: NNE, WNW, and NW (Fig. 44).

A2: DEPOSITIONAL AND TECTONIC MODELS FOR UPPER DEVONIAN SANDSTONES IN WESTERN NEW YORK STATE
Saturday Field Trip A2
Sandstone Outcrops from Allegany, Cattaraugus and Wyoming counties.
Guidebook for the field trip held October 7th, 2006 in
conjunction with the 35th Eastern Section AAPG Meeting and 78th
NYSGA Field trips held in Buffalo, New York
Gerald J. Smith and Robert D. Jacobi,
Field trip leaders and guidebook authors
UB Rock Fracture Group
Department of Geology
University at Buffalo
Buffalo, New York
Field trip sponsored in part by NYSERDA (New York State Energy
Research and Development Authority)
October 7th, 2006
We thank STATEMAP, EDMAP, NYSERDA, and Industrial Associates for
funding the research reported herein.
INTRODUCTION
Goals and Objectives
The intention of this field trip is to visit examples of the major clastic depositional environments prevalent in the Late Devonian Appalachian foreland basin. Additionally, we will examine how syndepositional faulting influenced, and in some cases controlled, the deposition of the major petroleum reservoirs. This field trip represents a continuing evolution in our understanding of the Upper Devonian depositional controls. Our earliest sandstone field trip (Jacobi et al., 1990) proposed turbiditic and storm depositional environments in a basin setting, consistent with the views of the time that the Appalachian Basin in New York State was essentially structurally featureless (except for rare folds in the south and the Clarendon-Linden Fault), and that the shorelines had been muddy. Our next field trip (Jacobi and Smith, 1999) showed that post-depositional faulting was likely, that sandy/silty shoreface environments were possible farther east, and that syndepositional faulting affected the Rushford Formation.
In this field trip, we will visit outcrops representing the major sandstone depositional systems (turbidites, shelf-ridges, beaches, and fluvial) which correlate with the main Upper Devonian oil and gas reservoir sandstones of the Elk, Bradford and Venango groups. At each stop we will point out sedimentary structures, trace fossils and bedding relationships that assist in determining the depositional environment. We will also discuss the significance of syndepositional deformation structures that occur ubiquitously within Allegany and Cattaraugus counties ofNew York State. These deformation structures, as well as other data sets (e.g., paleoflow indicators, structure contour maps, isopach maps, and ichnofacies), indicate that the depositional systems were influenced by an interplay among fault block rotation, eustatic sea level and sediment supply.
OUTCROP EXAMPLES OBSERVED ON THE FIELDTRIP
Our field trip will examine outcrop examples of the depositional environments that represent the major clastic reservoirs in the Upper Devonian that correlate with the Elk, Bradford and Venango reservoir sands to the south. Additionally, the fieldtrip shows examples of syndepositional faulting and the faulting influence/control on depositional environments. To cover as much ground as possible in the time allotted, we have chosen outcrops near roads, although these outcrops do not necessarily represent the best outcrop possible. Several stops are located on state property, or along roadcuts, and do not require obtaining landowners permission (although entry to Letchworth State Park will require an admission during certain periods of the year). Two stops are located on private property and will require obtaining permission from the landowner prior to visiting. Most landowners we have encountered are reasonable and offer no objection, unless they regularly discover trespassers or signs of trespasser. Please respect the wishes of the landowner.

A2: DEPOSITIONAL AND TECTONIC MODELS FOR UPPER DEVONIAN SANDSTONES IN WESTERN NEW YORK STATE

STOP 1. Letchworth Upper Falls Trail
Letchworth State Park contains a remarkable exposure of the Frasnian stratigraphic section along the 400 foot high cliffs and waterfalls carved by the Genesee River. Three major waterfalls are viewable in the park, the Lower, Middle and Upper Falls. Whereas much of the cliff sections are comprised of Gardeau Fm. shales and silty shales interbedded with thin turbidite/storm sandstones, the stratigraphic section above the Upper Falls exposes outcrops of the Nunda Formation (Clarke, 1897; Figure 23). The Nunda Formation is correlative to Elk reservoir sands, and is a good example of a thick turbidite sandstone TA8 ; however dewatering structures and wave-modification provide more variability than the original Bouma model (See figure 12).
The outcrop we will examine is exposed along the trail toward the Upper Falls and the High Bridge. The path and stairs were cut through 24 m of stratigraphic section, mostly fine-grained sandstone. In addition to examples of primary bedforms and sedimentary structures typical of turbidites, soft-sediment deformation and liquefaction features including sandstone dikes, roll ups, load casts, and zones of liquefaction with sharp boundaries are quite common in both the sandstones and the interbedded shales. Small down-on-the-east stratigraphic offsets ( -6 em) along northwest-trending (317} fractures typically coincide with observable deformation zones in the outcrop.
Bedding within some of the thicker sandstones appears similar to swaley-cross stratification, which suggests storm-wave modification/influence. Storm influence within the Nunda Formation is consistent with observations made higher in the stratigraphic section where "deep" water depositional environments (black shales and turbidites) show numerous bedforms (such as hummocky cross-stratification, 3-D ripples, linguoid ripples, e.g., Smith and Jacobi, 2002) that indicate depositional depth was within storm-wave base (Myrow et al., 2002; Pattison, 2005).
The specific depositional environment of the thick sandstones of the Nunda exposed at Letchworth is ambiguous. The thick nature of the sandstone beds, the striations on the base of the sandstone beds, the massive to planar (to possibly swaley bedding) all indicate proximal turbidites (TAB to T A-c). Such turbidites were typically are thought to represent submarine channel deposits. However, the lack of coarse basal sediments, and the constant thickness of the thick beds exposed along the High Cliffs at Inspiration Point suggest that either the channels were very wide, or that these represent a non-point source, more on the order of a storm-generated wash with a shoreface source. To the east we have found an erosional channel wall, indicating a channel (or slump scar) (Jacobi et al, 1994), and to the west, we suggested relatively thick Nunda sand beds represented sand lobe at the end of a channel (Jacobi et al, 1994), so it is possible that these sands at Letchworth do represent wide channels (or storm-generated wash that funneled into channels downslope).

STOP 2. Mills-Mills - Wiscoy Creek exposure (Figures 24, 25).
Wiscoy Creek at Mill-Mills where Mills-Mills Rd. crosses Wiscoy Creek. This is the upstream end of a fairly continuous outcrop that extends down to the village of Wiscoy. Roadcuts on either side of the bridge contain weathered exposures of the Hume Fm. black shale (Pepper and deWitt, 1951). The basal contact of the Hume Fm. with the underlying Mills-Mills Fm. (Smith and Jacobi, 2000) occurs below the Mills-Mills dam. The sandstones near the contact have a distinct petroliferous odor.
The underlying Mills-Mills Fm. (Smith and Jacobi, 2000) is lowest, thick sandstone package in the Famennian stratigraphic section in Allegany County. The Mills-Mills Fm. can be separated into upper and lower sandstone packets. The lower sandstone packet is a thick (∼3 m), upward fining sequence of fine- to medium-grained sandstone interbedded with thin gray shales and siltstone. The lower sandstones have tabular cross-stratification and displays prominent climbing ripples. The lowermost sandstone is thick (20-40 cm) and fine-grained sand; the sand bed has an erosional base and is channelized in cross-section. The upper sandstone packet ranges from 1 to 2 m thick and consists of thick beds (40 to 60 cm) of medium sandstones. Flute casts are observed on the soles of sand beds. Between the two sandstone packets is an interbedded section of thin gray sandstones, siltstones and shales that change upsection from gray to black. We have interpreted the Mills-Mills Fm. as offshore to restricted circulation deposits (Figure 12), with the sandstone packets deposited as TAB to TAC turbidites with TC -starting turbidites in the upper packet. Many of the lower turbidites are consistent with proximal turbidites deposited in channels, whereas the TC - starting turbidites appear more like lobe fringe sands.
The contact between the Mills-Mills Fm. and the underlying South Wales Fm. (Pepper and deWitt, 1951) is located downstream at the waterfall near the old RG&E powerhouse. The caprock of the waterfall is the basal thick fine-grained sandstone bed of the Mills-Mills Fm. The exposure of the Mills-Mills continues upstream with the sandstone beds forming small cascades up to the dam at Mills-Mills. The interbedded shales and siltstones grade from gray to black at the top of the formation.
Farther downstream, additional outcrops of the South Wales Fm. and Dunkirk Fm. (Clarke, 1903), are exposed along the southern bank, and in the north-flowing tributaries. Correlation of marker beds, such as the basal sandstones of the Mills-Mills and South Wales formations) indicate possible uplift along the fault east of STOP 2 (Figure 26).

STOP 3. Hillcrest Rd. Roadcut (Figures 27, 28)
Thick laminated sandstone (>3 meters thick) is exposed on either side of the road, where Hillcrest Rd. begins to climb the hill. The western side of the road exhibits the sharp contact between the Rushford Fm. (Luther, 1902) and the Caneadea Fm. (Chadwick, 1933), whereas the fracture controlled exposure on the eastern side gives an excellent view of the planar laminations and small, white, cloudy quartz pebbles incorporated into the fine-grained sandstone matrix.
The exposure of the Lower Rushford Mbr. here is interpreted as three stacked shore faces. The upper shoreface (planar beds) and foreshore (planar laminations) facies of each shoreface is eroded by an overlying transgressive surface of erosion. The Lower Rushford Member is correlative to the Bradford 3rd reservoirs along the New YorkPennsylvania border (Fettke, 1939; Manspeizer 1963; Smith, 2006).

STOP 4. Lunch Stop - Caneadea Dam and Rushford Lake.
STOP 4 - Lunch Stop -Caneadea Dam and Rushford Lake.
Beyond the fence is Caneadea Dam and Rushford Lake. Old photographs of the construction of the Caneadea Dam in 1929 show the north wall of the gorge (the side we are on) devoid of trees, and show quarries of the Rushford sandstones. By the dam, and around the eastern end of Rushford Lake, is the type exposure of the Rushford Formation as defined by Luther (1902), although close, hands-on access to these outcrops can be difficult depending on the lake level. When the lake level is lowered in the Fall, along the south gorge wall of the dam can be seen a series of north-striking, east-dipping, closely-spaced step faults; each with minor offsets (on the order of a few centimeters; Jacobi and Fountain, 2002). These small faults are consistent with well-logs that indicate a down-on-the-west fault of the Clarendon-Linden Fault System (Jacobi and Fountain, 2002) just west of this location. Other indicators of the fault include: 1) the north-striking valley (topographic lineament) where well-logs indicate a fault (Jacobi and Fountain, 1993), and 2) surface stratigraphy that suggests about 15.25 meters of offset across the fault (from Hillcrest Rd. outcrop (STOP 3) to the lake-level outcrops along the eastern edge of Rushford Lake.

STOP 5. White Creek at the Little John Rd. crossing (Figures 29, 30)
White Creek contains some of the best exposures of Luther's (1902) Rushford Fm. Road cuts north of Little John Rd. display the vertical section; creek exposures show the three, stacked shoreface sequences as well as exhibiting syn- and post-depositional deformation.
North of the bridge is a large exposure of the three-shoreface sequences that comprise the Lower Rushford Mbr. of the Rushford Formation. Basically similar to the Lower Rushford Mbr outcrops we examined near Rushford Lake, the outcrop here is much thicker (∼8.5 m). The variation of thickness of the Lower Rushford Mbr. coincides with faults of the Clarendon-Linden Fault System (Figure 31 ). The thicker regions possess preserved transgressive lags on offshore sequences, whereas the thinner regions possess eroded upper-shoreface and foreshore zones.
At the northernmost exposure of the outcrop, the 1st shoreface sequence forms the lowest step in the small falls of the outcrop. The top bed is a fossiliferous, medium to coarse sandstone that typifies the transgressive lags that occur on top of the shoreface sequences in the lower sandstone packet. Overlying the 1st shoreface is a thick interbedded section that is well exposure on the east cliff exposure. Soft-sediment deformation, primarily ball-and-pillow structures are well exhibited in the sandstone beds beneath the 2nd shoreface sequence. The 2nd shoreface sequence displays planer laminated beds typical of the shoreface sequence, but also includes a sediment slide (debrite) with a large sandstone olistolith surrounded by a muddy debris flow with deformed sand clasts. The top of the 2nd shoreface sequence occurs near the main falls/cascade. The sandstone displays small dunes (amplitude- ∼0.5 m, wavelength ∼2.0 m) that are a classic example of symmetrical ripples with tuning forks. The 3rd sandstone sequence forms the upper part of the cascade. The most noticeable feature is the transgressive lag that caps the sequence and the falls. This lag deposit contains large clasts of white, cloudy quartz as well as numerous brachiopod shell fragments and large red silt clasts. The underlying sandstone contains Rhizocorallium, Arenicolites and Thalassinoides, typical of a Glossifungites firmground. Overlying the 3rd shoreface sequence is the thick interbedded sequence that separates the lower sandstone packet from the upper sandstone packet.
We propose that the stress associated with the Acadian Orogeny caused a reactivation of basement faults, including those of the Clarendon-Linden Fault System and the Iapetan opening faults. Motion on these faults resulted in a number of small fault blocks that affected the basin topography. The motion and rate of slip and rotation among the faults was not identical, with the result that minor topographic highs and lows within the basin were created by small fault block motion that would create variable accommodation space over the faulted area (Figure 32).
We first proposed fault control of Upper Devonian sandstones to explain the occurrence - and preservation of shoreface sandstones in the Lower Rushford Member (Smith and Jacobi, 1998, 2000, 2001, 2003 and 2005; Jacobi and Smith, 2004). Syndepositional fault activity would enable coarser sediments to accumulate along active fault blocks at and adjacent to the intersection of the north-south trending Clarendon-Linden Fault System, and the northeast trending shoreline would act as a groin (Figure 33).
Upstream from the waterfalls located north of the bridge, deformation in the shales, silty shales and siltstones are readily observed; both deformed sandstones (ball and pillow) and pencil cleavages occur. At 0.25 mi (0.4 km) south of Little John Bridge, a repetition of the section appears to outcrop in White Creek. The upper part of the 2nd shoreface sequence and all of the 3rd shoreface sequence outcrops again at a small waterfall/cascade. Unlike the outcrop to the north of the bridge at Little John Rd., this outcrop does not have the thick transgressive lag deposits. However, a thin coquinite is observed at the top of the 3rd shoreface sequence and on top of the 2nd shoreface sequence, large cloudy quartz pebbles can be observed. The offset between the outcrop by the bridge and the repeated(?) section is approximately 11.6 m (∼38 ft), down on the west (or north).

STOP 6. Farnsworth Road - road cut (Figures 34, 35).
The roadcuts along the northern side of Farnsworth Road are exposures of the Hinsdale Fm. (Chadwick, 1933), (possibly a correlative to the "Pink Rock", an old drillers name), one of the stratigraphically higher Bradford reservoir sandstones. The Hinsdale Formation is comprised of thick sandstone beds that can be roughly separated into three packets. The thick sandstone lenses are typically fine-grained sandstone with thick lenses of coquinite that typically form basal lags, but also occur as small internal lenses (Figure 35). Thickness variability within the sandstones is high, further compounded by variable erosion of lower units by overlying sandstones.
Sedimentary structures are typically SCS with less frequent occurrences of TCS and tabular cross-beds. Small amplitude (and wavelength) symmetrical and interference ripples are commonly observed on the inner surface of large troughs and swales, implying complex, combined-flow during time of deposition (Figure 17). Thinner sandstones within the interbedded sections contain both asymmetric and 3-dimensional ripples and HCS; several thin sandstones also contain grooves and striations on the sole of the beds. Both the sandstones of the interbedded sections and the thick sandstone packets contain an abundant, but still low-diversity Skolithos ichnofacies. High frequency of Arenicolites and Skolithos occur in both the thinner sandstones as well as in the thick sandstone packets. A poorly developed Cruziana ichnofacies consisting of sporadic occurrences of Planolites and Teichichnus is found in the thicker interbedded sections.
We interpret the Hinsdale Formation as being deposited well above storm wave base, but still below fair weather wave base, most likely upper offshore to lower shoreface. The dominance of SCS and typical tempestite packets imply the Hinsdale Formation was deposited primarily by storms; the large amounts of fossils and the singular occurrence of red shales and sandstones suggest the shoreline was moving westward and terrestrial sediments were now being incorporated into the marine sedimentary deposits. The interpreted depositional environment is likely to be either nearshore storm-generated sand ridges or a storm-influenced barrier bars (Miall, 2000)
The roadcut displays the sharp-based sandstones that have lenses of coquinite throughout, but generally a thicker lag deposit at the base. Bedding in the sandstone is distinctly swaley, with large erosional surfaces cutting down into underlying units. At the lower roadcut, above the prominent, thick sandstone is a red to purplish red shale which aids the identification of the formation. The depositional environment for the Hinsdale and many of the similar thick sandstones between the Lower Rushford Member (STOPS 3, 4 and 5) and the Hinsdale Fm. are probably storm shelf ridges to storm-dominated barrier bars. The lateral variability observed at this location and many similar outcrops may explain the "lensing" nature of the Upper Devonian reservoir sands.

STOP 7. Little Rock City (Figures 36, 37).
Little Rock City is a State Forest under the supervision of the NYS Department of Environmental Conservation. The Salamanca Conglomerate (Carll, 1880) marks the basal contact of the Conewango Group with the underlying Conneaut Group. Sands and coarser units of the Conewango Group are correlative to the Venango reservoir sands in Pennsylvania. At this site exposures of thick orthoconglomerate (between 2 and 12 m thick) occur along the northern and eastern sides of the hill. House-size blocks separated along joints by periglacial expansion are abundant throughout the area; at least one was redeposited upside-down on top of another block by glacial activity.
Lithologically the Salamanca is an oligomictic quartz orthoconglomerate. The conglomerate consists of a medium to coarse sand matrix with clasts of cloudy white quartz. The clasts are oblate discoids that range from 0.5 centimeters up to 6+ centimeters in diameter. The oblate discoid shapes of the pebbles are typical of a wave-dominated shoreface. Large downlapping, graded cross-sets (1 to 4 meters high) and channels occur throughout the lower section of the outcrop. These bedforms may represent alluvial gravel bars. The form of the large cross sets is also similar to that found in estuarine tidal sand bars, but the thickness of the sets is high for the low (about 2 m) tidal fluctuations thought to characterize the Acadian Catskill Sea (e.g., Woodrow and Isley, 1983). However, the coarseness of the conglomerate indicates a high energy system, consistent with a tidal prism or bore in an estuary (and also consistent with the tidal gravels in the English Channel) that could have had a higher tidal range than the assumed 2 m Catskill Sea tidal range. Thus, these large crossbeds could be fluvial bars or tidal. Tabular crosssets and crossbeds with a herringbone pattern are observed above the large cross sets (although some of the herringbone is a "pseudo-herringbone" that results from trough cross sets in opposite directions). Both herring bone patterns, however, are consistent with reversing currents of tidal flux, as in a tidal channel. Overall, the large scale crossbeds, trough cross-sets, and the herringbone pattern suggest an estuarine/estuarine mouth depositional environment (Dalrymple, 1992) with a strong tidal influence. Paleoflow indicators from sedimentary structures are few, ripples and troughs on the upper surface occur sporadically.
We recognized several distinct facies within the Salamanca Conglomerate: facies A is the caprock of the Salamanca Conglomerate with the largest quartz clasts (5 cm). The conglomerate is generally disorganized with clasts orientated obliquely and perpendicular to bedding, except for the top surface; large carbonized wood fragments also occur. Facies B is a less resistant, more friable unit that exhibits herringbone bedding marking current reversals. Facies C is comprised of two thick, graded foreset packets each ranging 1 to 4 meters thick with large clasts accumulating along bedding planes with some asymmetric ripples. Facies D is a less resistant, more friable unit that contains fewer clasts, but generally does not exhibit reversals in bedding (Figure 47).
We have observed similar arrangement of facies at conglomerate blocks loose in Allegany State Park and adjacent areas, and at the "Bear Caves" area outcrops at Mount Seneca in Allegany State Park. The outcrops and loose blocks examined south of Little Rock City were generally sandier than the Little Rock City conglomerates, with clasts occurring only along bedding planes, and the clasts typically are not as large (maximum clast- 2 cm) as the clasts measured at Little Rock City.

A3: NORTHERN APPALACHIAN BASIN OIL AND GAS: HISTORICAL PERSPECTIVE
Arthur M. Van Tyne, Field Trip Leader.

A3: NORTHERN APPALACHIAN BASIN OIL AND GAS: HISTORICAL PERSPECTIVE

STOP 1. Bradford Oil Refinery
Our trip will proceed first to Bradford, Pennsylvania where we will visit an oil refinery owned by the American Refining Group. This refinery was started in 1881 and was formerly owned by the Kendall Refining Corp. This summer its 125 anniversary was celebrated by American Refining. It is now said to be, "the oldest continuously operating refinery in the U.S. and the oldest refinery in the world dedicated solely to processing 100% Pennsylvania Grade crude oil." It is one of only two small refineries still operating in all of central and western Pennsylvania. The Kendall company became famous for its two fingers in the air ad signifying that by the use of Kendall Pennsylvania grade oil you only needed to change the oil in your car motor every 2,000 miles instead of sooner with the use of other motor oils!

STOP 2. Rock City
After leaving the refinery we will proceed northeastward through Bradford and the Bradford oil field up into a hilly area between Bradford and Olean, New York. Just across the NY-PA State line we will stop at the tiny hamlet of Rock City. This place received its name from the huge rock formations which occur here. These rocks are the Olean conglomerate and sandstones of the Pennsylvanian Pottsville Group. These have been described as fluvial deposits by some researchers. Coming up the hill we passed outcrops of the underlying shales and sandstones of the Mississippian Pocono Group. At this locale the Olean conglomerate has been eroded into a literal "Rock City" and you can now walk down the enlarged joints as "streets" between the rock "buildings". Rock City is actually the technical geologic term for such erosional features. Because of the high purity of the silica in the Olean conglomerate nearby outcrops were quarried extensively during WWII for their silica content. Those pebbles were used in the manufacture of high fidelity radio frequency crystals which were installed in military and aircraft radios.
This rock city has been a Sunday outing site for people from the surrounding area since at least the early 1900s. The current owners, Cindy and Dale Smith, purchased the park from its former long-time owner only a few years ago. They have made a number of improvements to the trail system and have built a new pavilion with tables for picnickers which we may use for our lunch time today.

STOP 3. Pioneer Oil Museum
Our next stop will be at the Pioneer Oil Museum in the Village of Bolivar, NY. The museum was started in the 1970s by several local oil producers in an effort to save as much as possible of the artifacts of the early oil industry of this area. Old oil field equipment, data, maps and other information was usually discarded as various operations closed down. The museum has been quite successful in obtaining many gifts of old equipment and records from families and acquaintances of pioneer oilmen.
The building in which the Museum is now located was originally the McEwen Brothers Oilfield Supply Store. It was purposely located here in the heart of the Allegany oil field. An adjacent addition to the north was recently erected with the help of a New York State grant but the Museum has now outgrown its present site. A larger, and better constructed, building nearby has recently become available and the Museum will be moving to that site as soon as possible. That building, the old Hahn & Schaffner Supply Co. establishment, has been purchased by the New York State Oil Producers Association for the Museum.

STOP 4. Vosburg Lease
After leaving the Pioneer Oil Museum we will proceed eastward on Route 417 to the Vosburg oil lease owned by Plants & Goodwin. Mr. Paul Plants, the owner, is currently the President of the New York State Oil Producers Association and also of the board of the Pioneer Oil Museum.
The lease consists of 95 oil wells and 102 water input wells on a 275 acre tract. It has been operated for more than 60 years. This is a fully developed secondary oil recovery waterflood operation. The last major development took place in 1982 but one new oil well was drilled in the late 1990s. The oil wells are producing from the Richburg sandstone which is quite thick here. The average depth of these wells is 1,300 to 1,400 feet. About 2.5 million barrels of oil has been recovered from this lease over a long period of time. We will be walking in about 600 to 700 feet to visit a central power operation with a large eccentric shaped band wheel which pumps several wells with one rotation. This central power is one of only a few still operating in the New York-Pennsylvania oil fields

STOP 5. Gordon Brook Oriskany Gas Field.
We are stopping by here briefly only because it is on the way to our last stop. We are looking at the No. 1 Ramsey, the discovery well of the Gordon Brook Gas Field. The well was originally drilled in 1970 by Professional Petroleum Exploration but is now owned by Vandermark Exploration.
Gas Production is from the Oriskany sandstone at a depth of 3,840 feet. The well came in with a flow of 6 to 7 million cubic feet of gas per day. It was still producing gas up to a few years ago but now makes too much salt water when production is attempted. The well has produced about one-half billion cubic feet of gas. Just beyond the well head can be seen the separator where the brine is separated from the gas stream. The brine is then stored in the nearby tank until it can be disposed of properly. This is an example of Oriskany sandstone gas production from a small structure which is extensively faulted and where the Oriskany is pinching out. Gas has migrated into an offset in the pinchout line and later into the highly broken up sandstone to high areas along the faults. The field itself has produced about five billion cubic feet of gas and still has three producing wells.

STOP 6. Seneca Oil Spring.
This locale, sometimes known as the Cuba Oil Spring, is located about one and one half miles northwest of the Village of Cuba in central- western Allegany County, NY. This once active oil seep, well known to the local Native Americans, has been reported by many historians of the petroleum industry to have been, in 1627, the site of the first observation by Europeans of petroleum on the North American continent. This was said to have occurred when a Franciscan Recollect Friar, Father Joseph de la Roche Daillon, was taken to see the spring by a group of friendly Native Americans.
Fr. Daillon had been sent to Quebec, Canada by his superiors in France to search for and assist Fr. Nicolas Viel who was working as a missionary to the Hurons. He arrived in Quebec in May of 1625 and soon heard of the death of Fr. Viel in June of 1625. In mid 1626 he journeyed westward to the land of the Hurons. Later, in October of 1626, he journeyed westward again to the land of the Neutrals. While there, on July 18, 1627, Fr. Daillon wrote a letter to a friend in France telling of his experiences with Native Americans. That letter was published in 1636 by Fr. Gabriel Sagard Theodat in his "Histoire du Canada" and became the basis for later interpreters to claim that Fr. Daillon had been taken to see the Cuba Oil Spring.
The Cuba Oil Spring may still be the first place where Europeans saw petroleum in North America but, in the light of further research, we will have to change the date a bit. In a 1962 AAPG Bullet in the late Dr. John Wells, a distinguished Professor of Geology at Cornell University, published a paper in which he disputed the evidence for the Fr. Daillon visit. He quotes the above-mentioned Daillon letter as telling of the abundance of vegetables which were available as well as, "some very good oil" (huile). Wells felt that this term refers to an edible oil and believed that Fr. Daillon would have used the word petrole (rock oil) for such an occurrence as the Cuba oil spring because that term had already been applied to oil seeps which had been known in Europe and the middle east for hundreds of years.
He states that the earliest true reference to the Cuba oil spring was made in 1656 by later Jesuit missionaries who described it thusly, "As one approaches nearer to the country of the Cats (Eries) one finds heavy and thick water, which ignites like brandy, and boils up in bubbles of flame when fire is applied to it." This has also been documented in a paper by Thomas, et al., given at the Forum on History of Petroleum Geology at the May 1998 AAPG Annual Convention held in Salt Lake City, Utah.
The oil probably came from a subcrop of the Bradford First, and possibly also the underlying Chipmunk sandstone zone under 100 feet or so of valley fill . There is no crude oil visible at the Cuba oil spring now and there has not been any for many years. It is probable that the drilling of several shallow wells around the spring in the early 1860s depleted whatever was left of oil which had accumulated under the fill.
By the Treaty of Big Tree in 1797 one square mile around the spring was set aside as an Indian Reservation. The large plaque placed on the boulder of Olean conglomerate and dedicated in 1927 reads as follows:
SENECA OIL SPRING
Its history forms the first chapter in the development of the petroleum industry in America, a gigantic world enterprise transforming modern life.
1627 - Oil on American continent first recorded in this region by the Franciscan Fr. Joseph de la Roche Daillon.
1656 - Spring mentioned by Jesuit Fr. Paul LeJeune.
1721 - Visit by Joncaire.
1767 - Oil from this spring sent to Sir Wm. Johnson as a cure for his wounds.
1797 - Spring permanently reserved by Indians in Treaty of Big Tree.
1833 - Description of spring by Prof. Benjamin Silliman of Yale University.
Erected in 1927 by the University of the State of New York and the New York State Oil Producers Association.

A4: MIDDLE- UPPER DEVONIAN DEPOSITIONAL AND BIOTIC EVENTS IN WESTERN NEW YORK
Gordon C. Baird
Dept. of Geosciences
SUNY-Fredonia
Fredonia, NY 14063
D. Jeffrey Over
Dept. of Geological Sciences
SUNY-Geneseo
Geneseo, NY 14454
William T. Kirchgasser
Dept. of Geology
SUNY-Potsdam
Potsdam, NY 13676
Carlton E. Brett,
Univ. of Cincinnati
500 Geology/Physics Bldg
Cincinnati, OH 45221
INTRODUCTION
The Middle and Late Devonian succession in the Buffalo area includes numerous dark gray and black shale units recording dysoxic to near anoxic marine substrate conditions near the northern margin of the subsiding Appalachian foreland basin. Contrary to common perception, this basin was often not stagnant; evidence of current activity and episodic oxygenation events are characteristic of many units. In fact, lag deposits of detrital pyrite roofed by black shale, erosional runnels, and cross stratified deposits of tractional styliolinid grainstone present a counter intuitive image of episodic, moderate to high energy events within the basin. We will discuss current-generated features observed at field stops in the context of proposed models for their genesis, and we will also examine several key Late Devonian bioevents recorded in the Upper Devonian stratigraphic succession. In particular, two stops will showcase strata associated with key Late Devonian extinction events including the Frasnian-Famennian global crisis. Key discoveries made in the preparation of this field trip publication, not recorded in earlier literature, include: recognition of wave?-generated bedding and substrate channelization in styliolinid grainstone deposits of the Genundewa Limestone; discovery of a horizon in the lower Rhinestreet Shale with the chronozonally important goniatite cephalopod Naplesites and possibly dateable conodonts a few meters above a lag bed with Montagne Noire (MN) Zone 7 conodonts and a K-bentonite horizon presumably older than the Belpre Ash Bed of the southern Appalachian Basin (MN Zone 8); recognition of abundant phosphatic nodules in the Rhinestreet Shale cored by fish bones and scales; and discovery of a possible thin K-bentonite within the Pipe Creek Shale(= Upper Kellwasser Bed marking a global extinction-ecological disturbance event) allowing for potential absolute dating of that event.
This field trip differs from earlier ones in that it does not focus on any particular stratigraphic unit or problem. Rather we will provide somewhat of an ascending "Cook's tour" of units proceeding southwestward across Erie County from the lower medial Hamilton Group (STOP 1) to the Frasnian-Famennian mass extinction (STOP 7). This trip is, however, complimentary to our Sunday field trip (this volume) which focuses in more detail on the origin of the distinctive North Evans Limestone/Genundewa Limestone that we will see at STOP 2. For the benefit of participants, several of the newly-generated schematics are presented both in the Saturday and Sunday fieldtrip papers. In the tradition of NYSGA serving as a teaching forum, this fieldtrip is presented as "work in progress" and is designed to review controversies and ongoing questions. Although material covered is somewhat of an eclectic mix, this should lead to a broader array of issues to be addressed. Moreover, we have discovered several new features in several Late Devonian (Frasnian-Famennian) sections during field trip preparation that demand further investigation. These discoveries, discussed below, include the recognition of small-scale hummocky cross-stratification and substrate channelization in the Genundewa Limestone the finding of a new K-bentonite bed in the Pipe Creek Shale, the discovery of a geochronologically important occurrence of goniatites in the lower part of the Rhinestreet Formation, and the discovery of abundant phosphatic nodules containing fish bones and spines, also in the lower part of the Rhinestreet Formation. Thematic aspects of this trip will be: firstly, to examine evidence of current activity and bottom erosion in several basinal shale-dominated units intuitively viewed as "deep water" and "basinal," secondly, to look at distinctive facies and event horizons at different levels, and thirdly, to look at levels linked to key Late Devonian bioevents.

A4: MIDDLE- UPPER DEVONIAN DEPOSITIONAL AND BIOTIC EVENTS IN WESTERN NEW YORK

STOP 1. Charles E. Burchfield Art Museum and Nature Center
Town of West Seneca (see Figs. 4, 5).
When one of us (Baird) originally included this locality as a stop for a 1999 NYSGA fieldtrip (Baird et al., 1999), the present-day park was a vacant lot. The park has improved the access to this classic cutbank section along Buffalo Creek immediately upstream from Union Road. This stop is included, partly due to its relevance to our understanding of basinal submarine erosion processes, but also due to the fact that this stop was skipped owing to lack of time on the earlier fieldtrip. To reach the bank section we must wade across Buffalo Creek to the south-facing shale bank directly opposite the center. In the case of recent rains, the creek may be high, preventing our crossing. However, small, low exposures of the same channels are displayed at creek level on the south bank. These can be accessed in a high water scenario. At the bank, a prominent undulating contact is visible, separating gray to beige, calcareous mudstone below (lower division of the Levanna Member) from fissile, black shale above (medial part of Levanna Member; Figs. 4, 5). The undulating contact is an erosion surface which links eastward to a shell bed division of the Pompey Member of the Skaneateles Formation in central New York sections (see Baird et al., 2000). At this locality, conspicuous channel-like features are developed into the erosion surface which trend north-south, nearly perpendicular to the bank trend (Fig. 5). The channels range from a few decimeters to 1.5 meters width and up to 0.5 meters in depth. The dark shale fills in channels sometimes show smaller, nested channels within larger ones (Fig. 5c) indicating that channel formation was a continuous process during formation and burial of the contact. Parallel, north-south alignment of channels suggests a prevailing current direction during channel formation (Fig. 5b). Axes of channels sometimes contain concentrations of brachiopod and trilobite debris admixed with fish teeth and bones. The channels are interpreted by Baird and Brett (1991) to be ancient examples of submarine, erosional furrows, sensu Flood (1983); furrows form from sustained, unidirectional currents and can form in a variety of aqueous settings. However, they are found to be particularly common at modern continental slope breaks (Flood, 1983). We believe that the lenticular outcrop profiles of Leicester- and other detrital pyrite accumulations are the lateral perspective view of similar, debris-filled channels where surrounding mudrock has dewatered and flattened out channel relief.

Lake Erie visible to the right.
Low shale bluffs along shore are exposures of the Ledyard Member of the Ludlowville Formation (Middle Devonian).

Bridge over Eighteenmile Creek.
Upstream cutbank to the left shows the Middle Devonian Hamilton Group divisions unconformably overlain by Late Devonian strata of the Genesee Group.

Fishermen parking area
Fishermen parking area below pull-off to the left immediately south of the Lake Shore Road bridge over Eighteenmile Creek and small roadcut on Lake Shore Road to the right. This is STOP 1 of tomorrow's B3 Genesee Formation field trip (this volume).

Entrance to Frank Lloyd Wright Estate ("Graycliffs")
Entrance to Frank Lloyd Wright Estate ("Graycliffs") on the right. This newly restored attraction in the Buffalo area was a major access route to the shore cliffs along the lake in the 1980s before the great ladder below the estate fell into disrepair.

STOP 2. Lake Erie bluffs adjacent to the mouth of Pike Creek
Near Derby, Erie County (see Fig. 6). Proceed from end of long private driveway north of Pike Creek across yard to beach. Follow the beach southwest past Pike Creek to the beginning of a long bedrock cliff section bordering the beach. This section is in the southwestern part of a nearly continuous outcrop extending from Eighteenmile Creek almost to Sturgeon Point (STOP 4). Visible to the northeast are high cliffs exposing a long section, extending from the Wanakah Shale in the Hamilton Group, upward into the Frasnian Rhinestreet black shale. Units visible in the cliff south of Pike Creek, include: in ascending order, 2 meters of the Middle Devonian (Late Givetian) Windom Member, approximately 45 cm of lower Late Devonian (Frasnian) North Evans-Genundewa strata, 2.5 meters of gray and black shales of the West River Formation, and 2 meters of black Middlesex Formation. At this stop we will focus on the North Evans-Genundewa interval (Fig. 6).
This section shows the Windom Member-North Evans Limestone contact as a knife-sharp boundary separating chippy gray Windom shale rich in the small brachiopod Emanuella praeumbona, from dark, dysoxic beds of the Genesee Group (Fig. 6). The North Evans is anomalous in this area in that it is very thin and characterized by more detrital pyrite than typically found in sections closer to Buffalo. Furthermore, above the thin North Evans bed, and separating it from the overlying Genundewa Limestone, is a 25 cm-thick brownish black, flaggy shale unit that is unnamed (Fig. 6). This shale, containing flattened Styliolina, Pterochaeni, and wood debris, is very localized in distribution. At the North Evans type section on Eighteenmile Creek, adjacent to-, but immediately downstream from, the Amtrak railroad overpass, about four km east of STOP 2, the North Evans Limestone is a 6-12 cm thick lag accumulation dominantly composed of crinoidal debris, but also characterized by abundant conodonts, glauconitic grains, fish debris, and reworked concretions (Fig. 6). At the overpass bridge, the North Evans grades directly into styliolinid grainstone facies of the overlying Genundewa, and no dark shale is present. However, further downstream from that bridge to the northwest, a succession of creek bank sections shows dramatic thinning of the North Evans Limestone with the appearance of the feather edge of the dark shale as a nodular parting above the lag unit. Moreover, the North Evans undergoes a lateral, spectral change from a thick crinoidal unit to a thin layer of mixed carbonate grains and detrital pyrite that is distinctly more "Leicester"-like (Fig. 6). In the next key section (Fig. 6) along the Old Lakeshore Road (to be seen by participants of the Sunday fieldtrip B3: this volume), the intervening shale is 10-11 cm thick and the North Evans Limestone is only about 2-3 cm-thick. This intervening shale unit is herein interpreted as being a local basinal facies of the Genundewa Limestone in that its upper contact is conformable (non erosional) with the Genundewa Limestone. The westward North Evans allochem transition from dominantly calcareous to largely insoluble grains, accords closely to the appearance of the overlying dark shale, and it suggests that carbonate dissolution of the exposed lag was more intense in the more basinal subenvironment of STOP 2 than at localities further to the northeast.
At STOP 2, the North Evans Limestone rests on the Taghanic Composite Unconformity, a disconformity of large time-magnitude. Based on comparison with more continuous Devonian sections further east, the section at Lake Erie is severely truncated; approximately one third of the Windom Member, the entire overlying Tully Formation, and all of the succeeding Geneseo and Penn Yan formation successions are absent at STOP 2 (Fig. 2). The North Evans Limestone is a complex lag blanket famous for its fish material (Hussakoff and Bryant, 1918, Bryant, 1921, Turner, 1998) and conodonts (Hinde, 1879, Bryant, 1921, Huddle, 1974, 1981). North Evans conodonts are exceedingly abundant and diverse in species and types of elements. Elements of the North Evans fauna, both its fish debris and conodonts, can be traced from Erie County to the Genesee Valley. Preservation of the conodont elements is distinctive and varies from complete and pristine (light amber-colored) to broken, dark (almost black), and degraded (Color Alteration Index is 2 to 3). Colors of some of the fish debris are red, orange, gray and blue. The taphonomic history of the North Evans debris is complex and tests the limits of chronostratigraphic resolution (Kirchgasser and Koslowski, 1996; Kirchgasser and Vargo, 1998; Kirchgasser, 1994, 1998, 2001, 2002, 2004). The final taphonomic (burial age) of the North Evans material correlates to the youngest zone conodont in the mix, which is Ancyrodella recta of the upper part of Lower Frasnian MN Zone 2 (Kirchgasser, 1994); early whorls of the Lower Frasnian goniatite Koenenites (also MN 2 age) have been recovered from North Evans conodont residues at the type section at Eighteen mile Creek (Fig. 6a) and in lenses within the Genundewa Limestone at Linden in Genesee County (see Sunday illustrations). The North Evans, as with most lag beds, poses a depositional paradox; even though the lag content records an enormous span of time, the actual final depositional event producing the bed may have been geologically instantaneous.
The Genundewa Limestone at STOP 2 is a 30-40 cm-thick ledge composed of styliolinid grainstone-packstone carbonate (Fig. 6). It is typically brownish gray, massive to nodular, limestone which sometimes weathers into thinner, flaggy beds. Aside from Styliolina and the small bivalve Pterochaenia, fossils are scarce, usually small, and of low diversity. The goniatites Koenenites, Acanthoclymenia, and Tornoceras, as well as Manticoceras at the very top, occur in the Genundewa, but are rare or poorly preserved here; the conodont Ancyrodella recta of upper MN Zone 2 occurs in the North Evans Limestone and at the base of the Genundewa Limestone at Linden in Genesee County (see Sunday illustrations). As noted, the Genundewa is a pelagic limestone that probably represents oxygen-stressed, sediment-starved, basin slope conditions. However, at this locality, local erosional channelization is evident locally within the limestone (Fig. 6). These channels, as well as pervasive small-scale hummocky crossbedding and widespread alignment of Styliolina throughout the Genundewa, attest to significant current activity at the substrate, probably caused by deep-storm waves. The top of the Genundewa is gradational with the overlying shaley West River Formation; the topmost Genundewa becomes shaley and flaggy before giving way to the typical alternation of thin black and gray shale normally seen in the West River (Fig. 6). We herein interpret the interval from the sub-North Evans disconformity upward into the lower medial West River Formation is a Transgressive Systems Tract, starting with an erosional lowstand event recorded by the disconformity. The succeeding Genundewa is a condensed interval recording transgressive deepening within the basin and sediment-starvation on slope. A possible Maximum Flooding Surface is represented by a recurrent conodont-glauconite-bearing styliolinid bed within the basal West River Shale (Fig. 6); this conodont-rich bed has been traced westward from the Genesee Valley and is Huddle (1981) sample horizon 8122SD at Fall Brook, Geneseo (Kirchgasser et al., 1994, A-4, STOP I, fig. 10). Early highstand facies are represented by succeeding alternations of black and gray shale. This overall transgression probably represents a hybrid eustatic and flexural event within the basin.

STOP 3. Rhinestreet Formation at Lake Erie shore bluffs
Below at Seascape (private residence), 2 km northeast of Sturgeon Point near Derby, Erie County (see Fig. 7).
Near-vertical shore cliffs between STOPS 2 and 3 prevent our seeing units (Middlesex, Cashaqua formations) between the West River shale and the Rhinestreet shale. We will, however, see the top of the Cashaqua at STOP 5. STOPS 3-5 are directed to features in the lower part of the Rhinestreet Formation. Proceed across private yard to steps leading down the shore cliff. Turn right and proceed northward (down-section) along the beach to exposure of basal Rhinestreet strata.
The Rhinestreet Formation is one of the thickest Late Devonian black shale divisions in western New York. Hard, well-jointed, black shale makes up most of this section. Less organic-rich, recessive weathering dark gray shale intervals, as well as beds of greenish gray shale, can also be seen (Fig. 5). Conspicuous at several levels within the black shale intervals are large, often massive, septarial concretions. After walking northward on the beach, we will stop at a level displaying the largest concretions which occur about two meters below a 2.2 meter-thick gray shale unit (Fig. 7). The black shale facies within the Rhinestreet records intervals of severe dysoxia to near-anoxia within the Devonian basin associated with a broad time interval associated with global sea level highstand (Johnson et al., 1985). This organic-rich lithofacies accumulated in a relatively deep-water, stratified basin setting west of the prograding Catskill Delta in a foreland basin already maintained by collisional thrust loading (Ettensohn, 1998; see Fig. 1). Contemporaneity of Catskill Delta progradation is splendidly shown by the eastward splaying of Rhinestreet black shale divisions into the deltaic clastic wedge and by eastward passage of organic-rich basinal facies into shoreward, coarse, fossiliferous neritic facies and terrestrial red beds (Woodrow and Sevon, 1985; see deWitt, 1960; deWitt and Colton, 1959; Colton and deWitt, 1958; Sutton, 1963, and Kirchgasser et al., (1994) for evolution of correlations and unit terminology pertaining to eastern Rhinestreet and other West Falls Group subdivisions). The 2.3 meter to 2.4 meter-thick gray shale unit above the band of large concretions is the western distal "toe" of a progradational clastic pulse extending westward from the delta complex; this fine grained turbiditic or hemipelagic sediment was probably deposited during a sea level lowstand event which allowed prodelta muds to be exported far into the basin (Fig. 1). A turbiditic origin for part of this interval is suggested by the presence of a 2 - 5 cm thick siltstone bed at its basal contact at STOPS 3, 4, and 5 (Fig. 7); the basal surface (sole) of this layer displays erosional groove cast impressions and the top of the bed fines upward into featureless gray shale, suggestive of a turbiditic event. At STOP 3 (Seascape section), the basal lag of the siltstone bed yields conodonts of MN Zone 7 (Levin and Kirch gasser, 1994; Kirchgasser et al., 1994; Kirchgasser and Klapper, 1992). About one meter below the green-gray shale unit is a recessive notch that marks the position of a K-bentonite layer (Fig. 7). This altered ash bed is characterized by gray brown clays (kaolinite and mixed layered clays), bleached (pyroclastic?) micas, and secondary pyrite with minor quartz, calcite, plagioclase and ?apatite. The ash bed is graded with clay flakes in the upper part with even stronger parallel orientation than the clays in the overlying black shale. Lenses with the distinctive color and fabric of the ash occur intermittently for a few centimeters above the ash bed. As noted above, this ash bed is apparently not the Belpre Ash Bed of Tennessee (conodont MN Zone 8, Rotondo and Over, 2000) but an earlier event within the interval of MN Zone 7; the Belpre Ash has a radiometric date of 381.1 ± 3.3 Ma (Tucker et al., 1998; Kaufmann, 2006).

STOP 4. Rhinestreet Formation in Lake Erie shore bluff succession
Both to the northeast of-, and southwest of, the Sturgeon Point marina and car park complex near Derby, Erie County (see Fig. 7). We will, first, proceed on foot from the northeast end of the car park to the beach and follow the beach and continuous shale cliff section to a position near the outflow pipes from the Derby waterworks. The bluff section is nearly the same as that for STOP 3, except that most of the black shale interval below the gray shale unit is below lake level. The K-bentonite is not accessible here, but the gray shale unit is much easier to examine at this outcrop. Moreover, the black shale interval above the gray unit can be examined here (Fig. 7).
Within the gray shale unit in this locality are very thin, 0.2-6 cm-thick, black shale beds that often display sharp contacts and strong visual definition within the thicker gray succession. Buff gray concretions occur at two levels within the gray shale interval; these are concentrated closely below black shale bands and appear to be controlled by the presence of the bands. However, in a few places along the exposure, the thin, overlying black shale bands pinch out over the tops of the subjacent nodules. This suggests that concretion growth may have created differential paleorelief, perhaps due to early dewatering and differential settling of mud. The sharply defined thin, black bands pose an interesting question: Do they represent slow background deposition between turbiditic gray mud pulses, or do the thin black bands, themselves, represent some alternative type of rapid depositional event involving sedimentation, or resedimentation, of organic-rich sediment? Is it possible that the organic-rich sediment, instead, may have been originally pelletal, hence mobile and easily transported on the seafloor? This latter scenario, yet untested, could explain these sharp pinstripe bedforms as rapidly deposited, current-traction-generated features.
To the immediate northeast of the second waterworks outlet is a problematic structural displacement or offset at the level of the gray shale unit. At the lake edge, several fractures in the lower black shale division can be seen that are filled with gray shale that displays soft-sediment shearing and fracturing. Directly across from the area of shore fractures in the cliff face the gray shale unit thins to less than a third of its normal thickness across a distance of about 80 meters. However, debris on the beach conceals the intervening area and the full nature of this structure. The shore cliff and lake edge exposure displays excellent examples of joint networks, particularly for the black shale bands. These joints are believed by Gary Lash to have evolved during thermal-burial maturation of the black shale unit as the black shale units began to function as hydraulic top seals for moving fluids migrating up from below (see Lash, 2006; Lash and Blood, NYSGA Sunday trip B1, this volume). The different orientations of the various joint sets are believed to correlate to a series of far-field stress phases associated with the Allegheny orogeny (Lash and Blood, Sunday trip B1, this volume). This outcrop will be important as a stop on that Sunday trip.
Between 40 and 45 cm above the gray shale interval, within black shale facies, is a styliolinid-rich, hashy layer that is associated with widely-spaced, spheroidal concretions (Fig. 7). Close examination of the layer shows the presence of numerous, flattened goniatites (some partially pyritized), spotty concentrations of Styliolina, fish debris, large horizontal (arthropod?) trace fossils and scattered conodonts. The goniatites first identified during a survey of this section this past June, belong to the genus Naplesites in the family Beloceratidae. The lineage is characterized by compressed (discoidal), evolute shells with increasingly numerous, distinctly pointed lobes (and saddles) forming chevron-like patterns. The ancestor of the family is Probeloceras which in New York occurs below the Rhinestreet in the Cashaqua Shale (Sonyea Group; MN Zone 5); the group culminates with the extremely multi-lobed Beloceras, a genus still unknown in North America (Fig. 3). The discovery of the Naplesites horizon at Sturgeon Point (Stop 4) is important in that its presumed position in the lower Rhinestreet (House and Kirchgasser, 1993) is confirmed and the conodonts in the bed may prove to be datable. Naplesites is otherwise rare in New York and is known only from a few specimens (two species) described by Clarke (1898) far to the east from unspecified horizons and sections in the shales around Naples in Ontario County (Canandaigua Lake meridian). Probeloceras and Naplesites (as Mesobeloceras) are illustrated in Kirchgasser, Over and Woodrow, 1994, fig. 7).The close stratigraphic coincidence of useful conodonts, Naplesites, and the K-bentonite bed is important geochronologically and is the subject of ongoing work.
Return to vehicles and proceed by car around to the southwestern most parking area near the Sturgeon Point pier.
Proceed on foot to beach below car park and follow beach southwestward for approximately 120 meters to a low outcrop bench of black Rhinestreet shale at the lake edge.
This exposure of black shale is probably stratigraphically higher than the accessible beds northeast of the marina owing to regional dip effects, but the precise match of beds can not yet be made owing to the long covered interval between the two sections. Notable in this outcrop are several bands of concretions including a line of massive septarial concretions below the bench of very resistant black shale. At the top of the black shale bench are two closely-spaced horizons of small spherical to ellipsoidal, 0.3 - 4 cm-diameter, phosphatic nodules. Phosphatic nodules, well known from the upper part of the New Albany Shale and the Cleveland Shale, have not been reported from the Rhinestreet. The occurrence of nodules of this type has been interpreted as evidence of nutrient upwelling and high productivity in surface waters (see Robl and Barron, 1989). What is striking here is that many, if not most, of the nodules are nucleated by fish spines and scales. Small, sand-size grains within some of the nodules may be radiolarian tests.

STOP 5. Cashaqua Shale/Rhinestreet Shale contact
And overlying Rhinestreet strata exposed on Eighteenmile Creek and on access road leading to that creek, town of North Evans, Erie County (see Fig. 7). Enter access path opposite from North Evans Cemetery. Proceed on foot down sloped path to valley bottom. Cross flat area to edge of Eighteenmile Creek.
The Cashaqua Shale Formation of the Sonyea Group in southern Erie County is characteristically composed of gray fissile shale with a rhythmic succession of discoidal concretion bands (Buehler and Tesmer, 1963; Kirchgasser et al., 1997, fig. 9). The fauna of this unit consists of small mollusks and brachiopods, indicative of a dysoxic setting. Most notable, is the occurrence of numerous, often poorly preserved, goniatites belonging to the genera Manticoceras and Probeloceras which are flattened outside of the concretions and variably three dimensional inside of them (conodont Zone MN 5). Probeloceras is the ancestral genus of the family Beloceratidae and is followed by the genus Naplesites in the Rhinestreet Shale seen at Sturgeon Point, Lake Erie (Stop 3; see Kirchgasser, Over and Woodrow, 1994, fig. 7 for illustrations of Sonyea and West Falls Group goniatites). The prominent concretion layer with MN Zone 6 conodonts in the dark shales of the upper Cashaqua Shale is a septarian band with white or pink barite filling the shrinkage cracks and in places east of the Genesee Valley in Livingston County, replacing the shells of a rich molluscan fauna including the goniatites Manticoceras, Prochorites, Acanthoclymenia and Aulatornoceras; conodonts in the bed indicate MN Zone 6. The Prochorites is a species (P. alveolatus) known elsewhere only in Western Australia.
We will see the contact of this unit with the overlying Rhinestreet Formation of the West Falls Group in the bed of Eighteenmile Creek. The base of the Rhinestreet is sharp and marked by a seam of diagenetic pyrite which weathers to a rusty band in bank sections. No detrital pyrite has been found on the contact, but a rich association of conodonts is reported from the basal few centimeters of black Rhinestreet shale (Huddle, 1968; Kirchgasser and Klapper, 1992; Klapper et al., 1995).
Ascending the access path as we return to the vehicles, we pass the same lower Rhinestreet gray shale division that we examined at STOPS 3 and 4. Along the path this section is weathered and slumped, but, on the other side of Eighteenmile Creek, an excellent profile through this interval can be seen on a side tributary waterfall (Fig. 7).

STOP 6. Section of Pipe Creek Formation ("lower Kellwasser Bed")
And synjacent units at waterfall along south branch of Eighteenmile Creek, southeast of Eden, Erie County (see Figs. 8, 9). Proceed on foot from the vehicles to the lip of the large waterfall. The gray, silty, bioturbated unit in the falls face is the top part of the shaley Angola Formation of the West Falls Group. This is abruptly capped by the resistant, black, and well jointed Pipe Creek Formation of the Java group (Fig. 8) that forms the erosional lip of the waterfall. The 1.6 meter-thick Pipe Creek succession is followed upstream in the floor of the creek by softer, nodule-rich, gray and black shale facies of the Hanover Formation (Fig. 8).
The black Pipe Creek shale consists of massive, organic-rich facies that contrasts markedly with both the underlying Angola and overlying Hanover. The base of the Pipe Creek is abrupt on the Angola with the development of abundant diagenetic pyrite in the uppermost Angola. No erosional, detrital pyrite has been found at the contact, but a profound change in micro-texture is seen as one passes from the bioturbated, pyrite- suffused, uppermost Angola, into the laminated, black Pipe Creek (Fig. 9). As noted in the text, this unit has now been found to correlate to the "Upper Kellwasser Bed" which marks a major ecological reorganization and extinction event in global sections.
In the process of measuring this section for this field trip, we found an apparent K-bentonite bed 15 cm below the top of the Pipe Creek (Fig. 8, unit e). This layer consists of approximately 1 cm of soft clay rich in pyroclastic mica. The bed is currently being analyzed for its geochronostratigraphic potential (see text).
Above the Pipe Creek shale, there is an abrupt change to softer, gray shale that abounds in small concretionary nodules through an interval of about 1.5 meters (Fig. 8). One bed (Fig. 8, unit g), is packed with very small irregular nodules such that this layer somewhat resembles a muddy version of the classic condensed, gray, nodular limestone facies of European sections (Tucker, 1974; Tucker and Kendall, 1973). Fossils are common in the lower Hanover, but are usually very small. Diminutive gastropods, bivalves, and ostracodes make up most of the faunal trail mix. Goniatites, important as chronostratigraphic markers, occur sparingly in this interval at this section. However, at Java Village in Wyoming County, Delphiceras occurs just above the top of the Pipe Creek and somewhat higher in the section at Walnut Creek. At Silver Creek in Chautauqua County Sphaeomanticoceras and Crickites can be found, the latter genus being the highest zonal indicator of the Frasnian Stage (Fig. 3; House and Kirchgasser, 1993).

STOP 7. Upper part of Hanover Formation
Showing "Upper Kellwasser Bed" (Point Gratiot Bed) and basal contact of Dunkirk Formation on south branch of Eighteen mile Creek below New Oregon Road bridge, southeast of Eden, Erie County (see Figs. 10, 11). Exit vehicles and proceed to the downstream end of a long, west-facing cutbank along the south branch of Eighteenmile Creek below the New Oregon Road overpass. At this point gray shale of the upper middle part of the Hanover Formation are abruptly overlain by the black Point Gratiot Bed and succeeding gray and black shale beds of the upper part of the Hanover Formation (Fig. 10).
At this locality the Point Gratiot Bed is a black ledge that correlates to the "Upper Kellwasser Bed" of Late Devonian sections globally (see text). The topmost part of this ledge marks the actual Frasnian-Famennian extinction event of literature based on conodont work in western New York sections (see Over, 1997, 2002; Day and Over, 2002). Although the bed is composed of basinal black shale, fossils such as the probable anaptychus organ Spathiocaris, fish material, and wood debris can be found, particularly near the top of the bed. Moreover, elevated levels of Platinum Group elements occur at this level and other horizons in the boundary interval black shales (see Over et al., 1997).
Above the Point Gratiot Bed in this section is a 2.2 meter-thick upper Hanover succession of alternating gray and black shale beds that is capped by the black, well jointed Dunkirk Shale at the top of the bank (Fig. 10). The upper Hanover interval closely resembles the lower Rhinestreet gray shale unit owing to several discrete, thin, black shale bands which contrast sharply with the thicker gray lithology. Reworked pyritic burrow clasts and exhumed geopetally pyritized Tasmanites half-spheres (sensu Schieber and Baird, 2001) have been found in thin lags at the base of the thin black beds and at the base of the Dunkirk Shale. Immediately above the Point Gratiot Bed in this section is a thin, gray, bioturbated shale bed yielding small nodules (Fig. 10). Above this level, gray shale units become less visibly bioturbated and more non calcareous, suggestive of a long-term transgressive trend in the interval. Hanover deposits below the Point Gratiot Bed are markedly more calcareous, lighter colored, and more intensely bioturbated; upward changes across the Point Gratiot Bed probably reflect combined effects of transgressive deepening in the basin and adverse biological effects associated with the Frasnian-Famennian crisis. Participants should look for any sign of the fossiliferous "recovery bed" discovered by Jeff Over approximately one meter above the Point Gratiot Bed near Java Village (Day and Over, 2002). Although it is absent from many area sections, the STOP 7 outcrop has not yet been seriously checked for the presence or absence of this unit.

A5: LIFE ON THE EDGE: DEATH AND TRANSFIGURATION IN MUD
James J. Zambito
University at Buffalo, SUNY
University of Cincinnati
(zambitjj@email.uc.cdu)
Charles E. Mitchell
SUMMARY
The hypothesis of Coordinated Stasis posits that paleocommunities were relatively stable associations of species that tracked together through cyclical changes in their regional environment. When this tracking capacity was exceeded, they experienced relatively abrupt turnover in species composition, leading to a new interval of relative stability. This fieldtrip will focus on changes in community structure within the dysoxic, Ambocoelia-chonetid biofacies through the Wanakah and Lower Windom depositional cycles. We will also discuss the implications of the observed patterns for habitat tracking and community stability over longer time scales.
Recent work on the causes underlying patterns of biofacies distribution in time and space have led to a set of hypotheses referred to as Coordinated Stasis (Brett et al., 1996). Basic postulates of this model are that relative stability in the species composition of biofacies assemblages is a consequence of the ability of species communities to track preferred habitats and thereby avoid selection pressures that would otherwise disrupt community stability. When, on rare occasions, extraordinarily pervasive and abrupt environmental changes overwhelm the ability of species assemblages to track their preferred habitats, communities undergo a cascade of species turnover during which a substantial fraction of the formerly characteristic species go extinct and new species join the survivors to form a new and markedly altered community (Brett and Baird, 1995 and Brett et al., 1996). Habitat tracking is most easily generalized as an effect of sea level change. As sea level fluctuates through transgressive - regressive cycles, it follows that environments will synchronously fluctuate along an onshore - offshore gradient. If communities of organisms that share a preferred environment are to remain in this environment, then they will be forced to migrate up and down depositional slope through time with sea level. A critical issue for the testing of Coordinated Stasis, therefore, is the degree to which we can observe habitat tracking through the fossil record that is preserved within stratigraphic sequences.
The geographic setting of this trip is Western New York, with four outcrops from the Wanakah and Lower Windom Shale Members of the Hamilton Group, northern Appalachian Basin (Figures 1 and 2). The Middle Devonian Hamilton Group (Givetian Stage) is a dominantly fine-grained marine deposit. The sediments originated from the erosion of a mountain belt to the east that formed during the Acadian Orogeny. Most of the Hamilton Group records the interval between the second and third phases of mountain building that occurred from the collision of microcontinents into New England and the central Atlantic United States (Ettensohn, 1985; Ver Straeten and Brett, 1995); more specifically, the oblique convergence between the Laurentian and Avalon terranes (Ettensohn et al., 1988). This well-documented stratigraphic succession provides an ideal setting in which to examine the predictions of Coordinated Stasis. Indeed, it is the setting in which Brett and Baird (1995) developed the hypothesis.
Sea level change can occur at different magnitudes. Fifth- or sixth-order cycles are typically seen as 0.5-meter thick intervals in the study area. These consist of sub-symmetrical alternations of gray mudstone with concretionary, calcareous mudstones; an example of the latter is seen as the various trilobite beds of the Windom and Wanakah Shale Members. If these cycles were the product of Milankovitch forcing, they may have had a duration of ∼20 to 100 ka each (Brett and Baird, 1996). Fourth-order cycles commonly display a biofacies spectrum (described in more detail below) from a low diversity "Leiorhynchus" fauna, through the Ambocoelia-chonetid biofacies, to either a diverse brachiopod or coral-rich community (Brett and Baird, 1996) and, correspondingly, may have had an approximate duration of 0.1 to 1.0 m.y. Third-order cycles, representing 0.8 to 2.0 m.y., correspond fairly well to previously described formations in New York State, primarily because Cooper (1930, 1933) used the thin and widespread limestones in New York as the bases of formations. These limestones have subsequently been interpreted as third-order transgressive systems tracts (TST) by Brett and Baird (1996).
Both the Ludlowville and Moscow Formations, therefore, represent distinct third-order depositional sequences. On this trip we will examine the relationship between observed faunal changes and the early highstand systems tract (EHST) and highstand systems tract (HST) portions of these third-order cycles. Previous work suggests that both of these systems tracts are aggradational, whereas the late highstand systems tract (LHST, which we will not examine here) is dominantly progradational (Brett and Baird, 1996). These comparisons provide an opportunity to gage the degree of change in biofacies composition that accompanies sea level changes of various magnitudes. Unlike the case for the third-order cycles, fourth-order systems tracts have not been delineated for all of the intervals studied. Thus, fourth-order cycles are described here simply in terms of relative sea level change; that is, as fourth-order transgressive and regressive facies, rather than in terms of systems tracts. See Brett and Baird (1996) and Batt (1996) for full description of sequences and systems tracts addressed on this fieldtrip.
The Ambocoelia-chonetid biofacies is one of numerous recurrent fossil assemblages, or biofacies, described within the Hamilton Group of New York. These biofacies are representative of relatively narrow ranges of benthic environmental conditions (Figure 3). The Ambocoelia-chonetid biofacies represents conditions of low oxygen in a relatively quiet offshore setting below maximum storm wave base. Apparent habitat tracking through fourth-order sea level cycles results in symmetrical community replacement from "Leiorhynchus" to Ambocoelia-chonetid, to Athyris biofacies and vice versa (Figure 3). Ambocoelia umbonata, Crurispina nana, Devonochonetes scitulus, Sinochonetes lepidus, Longispina mucronata, Mucrospirifer mucronatas, Styliolina fissurella, Phacops rana, Greenops boothi and Auloporid corals are all commonly present in this biofacies along with numerous other, less common taxa (Figure 4). For a more detailed list of taxa found within this biofacies, see Zambito (2006).
Complicating any test of ecological stasis is the need to sample similar environments over time. This trip outlines the effects of sea level change, at the scale of both third- and fourth-order cycles, on organisms found within the dysoxic Ambocoelia-chonetid biofacies. Differences in community membership are described, and possible explanations for these differences are discussed. The primary application of this trip is a better understanding of considerations to be taken into account for studies that test communities for prolonged stability over time utilizing samples of from relatively small stratigraphic intervals.

A5: LIFE ON THE EDGE: DEATH AND TRANSFIGURATION IN MUD

STOP 1. Buffalo Creek
This stop focuses on the lowermost Windom strata, an Ambocoelia-rich gray shale. The base of the Windom Shale Member at this locality occurs directly above the Little Beards Creek Phosphate Bed which marks the disconformable contact of the Kashong and Windom Shale Members. This contact and the lower portion of the Ambocoelia-rich interval is accessible on the east-facing bank of Buffalo Creek almost immediately downstream of the new Bullis Road Bridge. The upper portion of the Ambocoelia-rich mudstones as well as the Smoke Creek Trilobite Bed are most easily accessed on the west-facing bank immediately upstream of the old Bullis Road Bridge.
The Ambocoelia-rich interval exposed at this locality represents the thickest exposure of this lower Windom unit on this trip, possibly representing a sub-basin (Gordon Baird, personal communication). The basal portion of these beds is a fissile and dark gray shale with only rare Ambocoelia umbonata, with a thickness of about I meter. This grades into a gray mudstone rich in both Ambocoelia umbonata as well as chonetids, including Longispina mucronata, Sinochonetes lepidus, Devonochonetes scitulus, and rare Devonochonetes coronatus. Near the top of the Ambocoelia-rich shales are at least six concretionary bands rich in styliolinids.
Ambocoelia umbonata becomes rarer above the second concretionary bed. The Smoke Creek Trilobite Bed attains a thickness of almost 1 meter at this locality and is underlain by approximately .5 meters of the Bay View Coral Bed.
The Ambocoelia-rich interval was deposited under regressive conditions, evidenced by the fissile dark gray shale below and the Bay View Coral Bed (requiring more oxygenated conditions) above. This interval is unique among those observed on this trip in that while it is representative of the Ambocoelia-chonetid biofacies, the absence of Crurispina nana and the presence of Longispina mucronata distinguishes this occurrence of the Ambocoelia-chonetid biofacies from the Wanakah Shale intervals described.

STOP 2. Cazenovia Creek
Proceed on foot back across US 20 Bridge over Cazenovia Creek Enter creek bed via driveway and trail on NW side of creek.
The lower Windom and upper Wanakah Shale Members outcrop on the north-facing bank of Cazenovia Creek, on either side of the US 20 overpass. Only the upper Wanakah is readily accessible at this locality, specifically the strata of and between the Spring Brook and Walden Cliffs concretionary intervals; occurring from the bed of the creek to approximately midway up the outcrop. This medium to dark gray shale unit contains abundant pyrite in the form of burrow steinkerns, pyritized fossils, and also pyritized shell hash horizons, which are commonly encased in concretions. A detailed description of these units, in particular the Spring Brook concretionary interval and its diagenetic history is described by Jordan (1968).
Unlike the lower Windom Shale observed at Stop 1, the Upper Wanakah shale interval contains abundant Crurispina nana. The intervals in which there is a presence of Leiorhynchus are indicative of a deeper water and lower oxygen biofacies, which corresponds to a slight decrease in the abundance of ambocoeliids.

STOP 3. Smoke Creek
The lowest strata exposed at this stop is the uppermost 1 meter of the Wanakah Shale Member. Overlying the Wanakah, where it has not been eroded, are pods of the Hills Gulch Bed of the basal Jaycox Shale Member. The cap of the falls at this locality is the Tichenor Limestone, a crinoidal grainstone containing many abraded fossil fragments indicative of a depositional depth near wave base (Brett and Baird, 1994). Immediately overlying this is the Ambocoelia-rich shales observed at Stop I. No styliolinid-rich concretionary beds are present at this locality, and specimens of Ambocoelia umbonata are typically deformed, presumably due to compactional stress. The Ambocoelia shale interval is overlain by the Bay View Coral Bed and the Smoke Creek Trilobite Bed. This stop is the type section of the Smoke Creek Trilobite Bed, a well-developed calcareous gray shale interval with abundant Phacops and Greenops. Similar to the interval exposed at Buffalo Creek, this occurrence of the Ambocoelia-chonetid biofacies lacks Crurispina nana and contains Longispina mucronata.

STOP 4. Lake Erie Bluffs
The exposure at the mouth of Eighteenmile Creek spans an almost complete Wanakah Shale section, with both the Tichenor Limestone and the basal Windom Shale Member visible when looking south along the shoreline. The interval examined, and easily accessible, at this section spans approximately 2 meters both above and below the concretionary trilobite beds of the lower Wanakah Shale Member.
Similar to the previous Wanakah Shale Member section examined at Cazenovia Creek, Longispina mucronata is absent from this intervaL Crurispina nana, while present, is extremely rare and has only been observed by the present authors in a few horizons (Figure 6).

A6: FAMOUS HAZARDOUS WASTE SITES AND FRACTURED ROCK HYDROLOGY OF THE NIAGARA FALLS AREA, NIAGARA COUNTY, NY
Guidebook for the Field Trip Held
October 7, 2006
in conjunction with the
78th Annual Meeting of the
New York State Geological Association
Buffalo, New York
Matthew W. Becker
Department of Geology
State University of New York
Introduction
More than 200 waste disposal sites have been identified within 5 kilometers of the Niagara River, which connects the largest freshwater system in the world (Yager, 1996). Chemical contaminants are likely to have leaked from one-third of these waste sites. Many of these hazardous waste facilities have discharged waste into the Lockport Group that underlies the region. The Silurian-aged Lockport Group (or Lockport Dolomite according to USGS designation) is fractured along highly continuous bedding planes that bear and conduct water. The hydrogeology of the Niagara Falls Area, and therefore the ground-water contamination, are dominated by fractured rock hydrology.
The objective of this field trip is to provide a first hand exposure to some of the most well-known ground-water contamination sites in the context of the regional hydrogeology of the Niagara Frontier. We will visit four Superfund sites, Love Canal, Bell Textron, 102 Street, and Hyde Park, and then view the stratigraphy of the Lockport group at the Niagara Gorge. Along the way, general principles of fractured rock hydrology will be discussed. Because this trip is held on a Saturday, we will not have direct access to any of the sites themselves, but we should be able to see enough from the road to provide the hydrogeologic setting.

A6: FAMOUS HAZARDOUS WASTE SITES AND FRACTURED ROCK HYDROLOGY OF THE NIAGARA FALLS AREA, NIAGARA COUNTY, NY

STOP 1. Bell Aerospace Textron Site
The Bell Aerospace Textron Site is located in Wheatfield, New York, adjacent to the east site of the City of Niagara Falls. In the 1950's and 1960's TCE (trichloromethane) was used in the production process. Waste TCE was discharged to a shallow "neutralization" pond which infiltrated to the Lockport Group. By 1990, a 280-acre aqueous plume in the Guelph Dolomite (Figure 3) that contained TCE and its metabolites (DCE, dichloroethane and VC, vinyl chloride) extended about 4,300 ft south of the pond. A 20- acre dense, non-aqueous phase (DNAPL) plume of TCE extended 620 ft south of the pond. A pump-and-treat remediation system consisting of six wells near the pond and five wells 2,900 ft down-gradient from the pond began operation in 1993 to decrease the size of the aqueous plume and prevent its further migration.
The study by Yager (2000) illustrates some of the interesting characteristics of contaminant transport in bedrock systems. Flow and transport is through individual fractures with hydraulically estimated apertures of 1 to 1.5 mm. The transmissivity estimated through model calibration is 140 m2/day. As is often the case in bedrock studies, it is difficult to estimate the effective porosity. Yager (Yager, 2000) modeled effective porosity between 0.3 and 3 percent. The smaller number was consistent with the hydraulic aperture estimates and the larger was more consistent with transport modeling. Ground water velocities are thought the be on the order of meters per day, but are difficult to confirm given the uncertainty in effective porosity. The plume (Figure 3) is nearly as wide as it is long, reflecting the similarity in transverse and longitudinal dispersion that is thought to be characteristic of transport in fractured bedrock. Another important consideration in transport is the exchange in contaminant mass between fractures and rock matrix through a process known as "matrix diffusion". Molecular diffusion becomes an important consideration in fate and transport through dual porosity media. Models of plume transport and biodegradation were extremely sensitive to the rate of matrix diffusion assumed. Matrix diffusion is difficult to measure independently for application to field studies.
The plume at the site appears to have reached a dynamic equilibrium between the rate of TCE dissolution and the rate of removal through pumping and biodegradation. The presence of dissolved mass in the rock matrix and non-aqueous phase pools in fractures provides a constant source of TCE. Degradation rates for TCE are 21 to 25 days, DCE are 170 to 230 days, and VC are 18 to 23 days. Consequently, if the source could be removed the plume would disperse rapidly. Removal of all the contamination source in the rock and fractured bedrock is considered impractical given current remediation technology. A summary of the site is provided in Appendix 81 (U.S. Environmental Protection Agency and New York State Department of Environmental Conservation, 2005).

STOP 2. Love Canal
It is unlikely that any reader has not at least heard of Love Canal. It is, perhaps, the most famous ground-water contamination site in the country and launched the creation of the Superfund program in the United States. A brief History is presented in Appendix B1. The site was removed from the EPA's National Priority Listing (Superfund List) in 2004 and its final remediation actions have been implemented. The hydrogeology underlying Love Canal consists of a shallow system of silts and fine sands, underlain by confining layers of lacustrine clays and glacial till, which are underlain by the Lockport Dolomite. Wastes were emplaced in the canal excavated as part of William T. Love's plan to create a canal that connected the Niagara River to Lake Ontario. The site resembled a geologic bathtub filled with waste and capped with soil. After wet periods, leachate would rise to the surface in swales where humans could come in direct contact with the material. Leachate also seeped into nearby basements.
The 1988 Superfund record of decision for cleanup of the site was "The selected remedial action for this site includes: excavation and solidification/stabilization of 7,500 yd3 of soil; placement of solidified soil back in excavated location; installation of a RCRA cap; ground water monitoring; and implementation of treatability studies for solidification process." This essentially is the solution that was ultimately implemented at the site. Occidental Chemical owns the site and operates the treatment facility housed on the property. Hazardous materials are stripped from the ground water and sent to an incineration facility in Texas for ultimate disposal.
The hydrogeology underlying Love Canal consists of a shallow system of silts and fine sands, underlain by confining layers of lacustrine clays and glacial till, which are underlain by the Lockport Dolomite. The sediments are believed to form a liner that prevented extensive contamination of the bedrock. Constant pumping within the excavated canal appears to have reversed the ground-water flow at the canal and removed dissolved contamination from both the bedrock and the overburden sediments.

STOP 3. OCC S-Area
The S-Area site is an eight-acre landfill on Occidental Chemical Corporation's (OCC) Buffalo Avenue Plant. The site is located approximately 200 yards north of the Niagara. The site was used primarily from 1947 to 1961 for the disposal of approximately 63,000 tons of organic and inorganic chemicals. Chemicals deposited at the site included chlorobenzenes, organic phosphates, acid chlorides, phenol tars, thionyl chloride, chlorendic acid, trichlorophenol, benzoyl chloride, liquid and chlorotoluene-based disulfides, metal chlorides, thiodan, and miscellaneous chlorinated hydrocarbons. The EPA Fact Sheet for this site is provided in Appendix B3.
The S-Area Landfill is historically significant because it was at this site that the term "non-aqueous-phase-liquid" (NAPL) was first used (Pankow and Cherry, 1996). The landfill is located immediately adjacent to the former City of Niagara Falls Water Treatment plant. The plant drew water from the Niagara River via a bedrock tunnel. Contaminants from the S-Area landfill leaked into this tunnel, contaminating water supplies. The entire treatment plant was abandoned and a new plant constructed in 1997. The site of the former water treatment is clearly visible between the S-Area and new treatment plant.
The contaminated ground water flowed toward the treatment plant intakes and the Niagara River prior to remediation. Ground-water flow was through three zones: (1) overburden sediments, (2) shallow weather bedrock, and (3) deeper bedrock (Figure 5). Due to the site's close proximity to the Niagara River, contaminated ground water discharged to the Niagara River or to bedrock beneath the river. The remediation strategy is containment. A combination of pumping, drains, and slurry walls are used to create an inward hydraulic gradient at the site. Effluent is treated onsite and then discharged to the Niagara River.
Unlike the Bell Aerospace Textron site, little is known about the natural attenuation of chlorinated solvents. This is essentially because the plume extends immediately to the Niagara River making water quality observation difficult.

STOP 4. Hooker (OCC) Hyde Park
Occidental Chemical Corporation's (OCC) Hyde Park site is a 15-acre landfill in northwest Niagara Falls, less than one-half mile from the Niagara River. From 1953 to 1975, the company (then Hooker Chemicals and Plastics) deposited approximately 80,000 tons of chemical wastes at the site. This is arguably the largest mass of DNAPL contamination site in the United States. The hazardous materials disposed on site included 3,300 tons of 2,4,5-trichlorophenol (TCP) wastes, which are known to contain significant amounts of 2,3,7,8-tetrachlorodibenzo-p-dioxin (TCDD); approximately 0.7 - 1.6 tons of dioxin are believed to be associated with the TCP. Chlorinated organic wastes, including hexachloropentadiene derivatives, chlorendic acid, chlorinated toluenes, benzenes and phenols, predominate at the site. The former drainage stream of the landfill, Bloody Run, which flows into the Niagara River, was historically contaminated with organic chemicals, including dioxin. A clay cap and a shallow leachate collection system were installed at the site in 1979. A summary of remediation activities can be found in Appendix B4.
The site is underlain by Pleistocene overburden deposits that overlie the Lockport bedrock. Overburden sediments are glacial till, lake deposits, and some localized sand and gravel deposits. The principal water bearing zones are in the Lockport Group are the weathered bedrock surface and horizontal- fracture zones that coincide with stratigraphic contacts (Yager, 1996). The underlying Rochester Shale is thought to have much lower hydraulic conductivity than the Lockport and therefore constitutes the lower limit of possible contamination below the Hyde Park Site (S.S. Papadopulos & Associates, 2001). Historically, contaminated ground water seeped at the Niagara Gorge face east of the power dam (Figure 6).
The Hyde Park Site is another excellent illustration of the difficulties faced when characterizing and remediating ground water in fractured bedrock. As in the case of Bell Textron and S-Area, DNAPL has seeped into the bedrock and cannot be removed entirely. The remediation strategy is, therefore, perpetual containment. At Hyde Park, containment is achieved by the placement of pumping wells that are designed to assure a constant inward hydraulic gradient at the site boundary. In bedrock, pumping is carried out at three depths that coincide with water bearing fractures identified by local testing and regional modeling (Figure 7). In spite of the use of 15 purge wells over three zones, it has been difficult to capture the dissolved phase plume. A recent modeling evaluation of purge well effectiveness showed that only the southern portion of the plume was being captured in the upper bedrock zone. Most of the plume was being captured, however, in the middle and lower zones.
Actually monitoring the ground water velocity vectors has proved challenging at Hyde Park. Monitoring wells were installed radially away from the center of the contaminated area with the intention of ensuring inward hydraulic gradients. Local heterogeneities in transmissivity make such monitoring difficult. The local changes in head due to local changes in transmissivity serve to confuse the hydraulic data. The actual ground-water divide induced by on-site pumping cannot be determined with any certainty. Although the MODFLOW model calibrated to monitoring data indicates that the most of the contaminant plume is currently captured (S.S. Papadopulos & Associates, 2001), this cannot be confirmed with head monitoring data.

STOP 5. NYPA Access Road
This road leads to the base of the NYPA dam where a fishing platform is located. An excellent section of the Lockport and Clinton Group is exposed. Recent installation of fencing has obscured the view somewhat. About halfway down the road an ephemeral seep is fenced to keep visitors away from contaminated ground water derived from Hyde Park. In spite of the hydraulic controls at Hyde Park, this seep is occasionally active. The revised stratigraphy of the Lockport and Upper Clinton are provided here (Figures 8 and 9).

STOP 6. NYPA Power Vista Center
We will stop at the Power Vista Visitor's Center (http://www.nypa.gov/vc/niagara.hhn) to get a view of the forebay canal. When the Niagara project produced its first power in 1961, it was the largest hydropower facility in the Western world at the time. The Niagara project, located about 4.5 miles downstream from the Falls, consists of two main facilities: the Robert Moses Niagara Power Plant, with 13 turbines, and the Lewiston Pump-Generating Plant, with 12 pump-turbines. In between the two plants is a forebay capable of holding about 740 million gallons of water; behind the Lewiston plant, a 1,900-acre reservoir holds additional water.
The excavation of the forebay into the Lockport provides a unique view of how water moves through fractured bedrock. Vegetation and water staining, along with ice formation in the winter, designate seeps along the canal walls. Notice that these seeps occur along specific bedding plane contacts, but only in specific locations. This "flow channeling" is characteristics of water flow through fractured bedrock. It implies that (1) effective porosity may be much smaller than would be estimated from fracture occurrence and (2) monitoring of contamination in bedrock may be partly a matter of luck. When effective porosity is overestimated, ground-water velocity is underestimated, based upon hydraulic arguments. This is due to the expectation in Darcy's Law that effective porosity (neff) relates specific discharge (q) and average linear velocity of a contaminant (v); v = q/neff.

STOP 7. Devils Hole State Park
If time allows we will hike down the trail that leaves from Devils Hole State Park. This trail affords an opportunity to seem more Lockport exposure up close as well as a natural cave. There are few examples of karst morphology in the Lockport Dolomite. It is unlikely that karst effects ground-water flow in the region. This cave was probably formed during the carving of the gorge by the Niagara River. It does provide an up close and personal look at the Lockport Dolomite.

B1: THE UPPER DEVONIAN RHINESTREET BLACK SHALE OF WESTERN NEW YORK STATE-EVOLUTION OF A HYDROCARBON SYSTEM
Gary G. Lash
Department of Geosciences
State University College
Fredonia, NY 14063 USA
and
DAVID R. BLOOD
Chesapeake Appalachia L.L.C.
Charleston, WV
Organic-rich, low permeability, sometimes fractured, shale deposits, once ignored by drillers seeking more readily understood plays and faster returns on their investments, are boosting the fortunes of mid-sized producers across the United States and becoming an increasingly important resource base. This burgeoning interest is driven largely by increased natural gas prices, improved completion technologies and a lack of conventional reservoir targets. However, there is no universal model applicable to each and every unconventional or continuous-type reservoir. Indeed, most vary in terms of basic stratigraphic facies distribution, mineralogy (i.e., quartz content and clay type and content), fracture parameters (length, orthogonal spacing, connectivity, anisotropy), porosity and permeability, and rock mechanical properties. There clearly is a need to further our understanding of these economically important deposits. This field trip focuses on the evolution of the heavily fractured Upper Devonian Rhinestreet black shale, which has served not only as a hydrocarbon source rock, but also as its own seal and reservoir, the essence of the unconventional reservoir. Specifically, we address the Rhine street shale in terms of the hydrocarbon system, which comprises three basic elements (Demaison and Huizinga, 1994). The first component, charge, is the hydrocarbon volume available for entrapment and is dependent upon source rock richness and volume. Migration, the second parameter, can be thought of in terms of the lateral movement of hydrocarbons via permeable carrier beds versus vertical migration via faults and, more to the point of the Rhinestreet shale, fractures. The final component of the petroleum system is entrapment, which involves consideration of those rock properties that would favor the arrest of migrating hydrocarbons.
We will consider the Rhinestreet shale from deposition through early diagenesis, including mechanical compaction, followed by the early and relatively shallow generation of abnormally high pore fluid pressures caused by a marked increase in sedimentation rate at the end of the Devonian. Soon after this, at the onset of the Alleghanian orogeny, the Appalachian Basin of western New York state and western Pennsylvania was uplifted resulting in the initial phase of fracturing in the Rhinestreet shale. This was followed by rapid burial of the organic-rich rocks to the oil window and the generation of natural hydraulic fractures (joints) in response to catagenesis. Indeed, the jointing history of these rocks provides a record of the complex overpressure history of the Rhinestreet shale and other black shale units of the western New York state region of the basin.
Results reported herein represent the record of six years of work on the Upper Devonian shale-dominated succession exposed along the Lake Erie shoreline and general vicinity, the Lake Erie District (LED). This work is now being extended into the subsurface of western New York state, northwest Pennsylvania, and northeast Ohio by use of wire line logs. During his involvement in the study of the Rhinestreet shale, DRB was supported by two SUNY Fredonia Summer Undergraduate Research Fellowships. Most recently, his work, which constitutes the better part of his MS thesis at the University of Buffalo, was supported by an AAPG grant-in-aid. We acknowledge Peter Bush and his staff at the University of Buffalo, South Campus Instrumentation Center, School of Dental Medicine, for their unflagging help with the electron microscopy, Ross Giese and Anja Dosen, University of Buffalo, for the XRD analyses, and Taury Smith, Rich Nyahay, and James Sandor of the Reservoir Characterization Group at The New York State Museum, Albany, New York. Finally, we thank Terry Engelder for many fruitful discussions on compaction, jointing, the geology of the Catskill Delta Complex, and more in the field, over the phone, and via E-mail.

B1: THE UPPER DEVONIAN RHINESTREET BLACK SHALE OF WESTERN NEW YORK STATE-EVOLUTION OF A HYDROCARBON SYSTEM

STOP 1. PIKE CREEK/LAKE ERIE SHORELINE
Significance:
I) the preferential generation of early NS-trending joints at the contacts of gray shale and black shale;
II) geographic variations in the magnitude of the pre-Upper Devonian or Taghanic unconformity and the significance of the so-called "Driller's Tully";
III) preferential propagation of NW and ENE joints in the Rhinestreet black shale.
I. Working down creek toward Lake Erie, the first exposure encountered reveals the contact of the West River gray shale and overlying Middlesex black shale (Fig. 40). A prominent NS-trending joint set as well as ENE-trending joints can be observed here (Fig. 41). The NS joints terminate at the contact or a short distance above it, a theme common to other Upper Devonian black shale units of western New York state as described earlier. We interpret the generation of the NS joints to be a response to the passage of a peripheral bulge produced by loading of the Laurentian plate edge at the onset of the Alleghanian orogeny across the western New York state region of the Appalachian Basin. Reduced compressive stress as a consequence of uplift-related thermoelastic loading initiated joints in higher modulus diagenetic carbonate. Propagation of these joints through the host gray shale was aided by moderately overpressured pore fluids at the top of the gray shale, perhaps sealed in by the overlying low permeability black shale.
II. The mouth of Pike Creek exposes the Middle-Upper Devonian interval from the Tichenor limestone (perhaps the top of the Ludlowville shale, depending upon the lake level) to roughly two-thirds the way up the Middlesex black shale (Fig. 40). This exposure provides the opportunity to (1) consider the magnitude of the Taghanic unconformity in the subsurface of western New York state and (2) offer some preliminary thoughts on the subsurface stratigraphy of the Moscow-Tully interval in this region of the basin. De Witt and Colton (1978) placed the unconformity at the contact of the Penn Yan shale of the Geneseo Formation and the North Evans limestone of the underlying Moscow shale, Hamilton Group (Fig. 40). Rocks absent this exposure include, in descending order, an unknown thickness of the Penn Yan shale and the underlying Geneseo black shale, the Tully limestone, and some thickness of the Moscow shale (Windom Member). The hiatal extent or magnitude of the Taghanic unconformity diminishes to the south and east as the Moscow shale increases in thickness and the Geneseo black shale and Tully limestone appear and thicken (Fig. 42). De Witt et al. ( 1993) maintained that the disconformity confined the Middle Devonian Geneseo black shale to the eastern region of the basin in New York state; however, Upper Devonian black shales, including the Dunkirk, Pipe Creek, and Rhinestreet shales were largely restricted to the western part of the basin by the unconformity. However, the distribution of the Middle Devonian Oatka Creek black shale of the Marcellus Formation (Hamilton Group) appears to have been unaffected by the Taghanic unconformity.
We are using wireline logs to define stratigraphic trends in the subsurface of western New York state, northwest Pennsylvania, and northeast Ohio. Figure 40 links the stratigraphy observed at the mouth of Pike Creek (as well as the cliff exposure to the northeast along the beach) to the Eckhardt-Agle 1 gas well (API# 17937) located ∼10 km southeast of this location (Fig. 42). The Middlesex shale and Tichenor limestone are easily recognized on the gamma-ray log. Locating the North Evans-Genundewa interval, and, therefore, the Taghanic unconformity is more difficult. The Genundewa limestone, ∼21 - 30 cm thick on Eighteenmile Creek between the beach exposure and the Eckhardt-Agle 1 well (Fig. 42), is difficult at best to be resolved on a normal gamma-ray log. Thus, our placement of the unconformity is based on extension of the subsurface stratigraphy from the east, where the Tully limestone and overlying Geneseo black shale are present, to the Eckhardt-Agle 1 well. There is a marked increase in the thickness of the West River shale to the south away from the lakeshore (Fig. 40). Cross-section 1 (Fig. 42) carries the interval exposed at the mouth of Pike Creek (i.e., the base of the Tichenor limestone to the base of the Middlesex shale) east-southeast into northeast Alleghany County. The Geneseo black shale first appears in Wyoming County followed in northwest Alleghany County by the Tully limestone. Notice also that the West River/Penn Yan interval displays an almost six-fold increase in thickness to the east. The Tichenor limestone also thickens to the east. In sum, cross-section 1 displays, over a relatively short distance, the westward onlap of progressively younger deposits over the Taghanic unconformity as well as the truncation of underlying units, both typical traits of the unconformity elsewhere in the Appalachian Basin (Hamilton- Smith, 1993).
An especially interesting unit that we are mapping in the subsurface, the so-called "Driller's Tully," is a carbonate interval that is less clean (i.e., 20 API units or more) than the true Tully limestone. There is a good amount of confusion regarding this unit; Oliver et al. (1969) placed the "Driller's Tully" of Erie County, Pennsylvania, and Chautauqua County, New York, well down the Hamilton Group, perhaps at the level of the Tichenor limestone. Similarly, Wright (1973) referred to the "Driller's Tully" as limestone "B" and suggested that the Tichenor limestone is a tongue of limestone " B." This interpretation was accepted by de Wltt et al. (1993) who also echoed Wright's proposition that the Tully limestone is not present in the subsurface of Chautauqua County and northwest Pennsylvania. De Witt et al. (1993) pointed out that conodonts collected from the upper part of the "Driller's Tully" in a core from the Eastern Gas Shales Project well PA3 in Erie County, Pennsylvania, included elements of Polygnathus linguiformis,
We agree with Rickard (1989) that the Tully limestone is present in the subsurface of Chautauqua County and northwest Pennsylvania. Indeed, our analysis of close-spaced wells in far western Chautauqua County indicates that Wright's (1973) Tichenor limestone and limestone "B" is actually the Tully limestone and underlying "Driller's Tully," the latter resting unconformably on the Tichenor limestone. However, we interpret the "Driller's Tully" to be a discrete unit in erosional contact with the Moscow shale, and locally the Tichenor limestone, rather than a stratigraphic equivalent of the Moscow shale and Tichenor limestone as suggested by Rickard (1989). Truncated calcareous intervals of the Moscow shale and locally abrupt changes in thickness of the "Driller's Tully" confirm the erosional contact of this unit with underlying deposits. The "Driller's Tully" is recognized in the subsurface of northern Allegany County, that area of Chautauqua County where the Tully is present (Fig. 43) and parts of Warren and Erie counties, Pennsylvania. The contact of the "Driller's Tully" and overlying Tully limestone is a thin shaley interval (Figs. 42 and 43). However, the local absence of this interval and as well as the occasional absence of the "Driller's Tully" beneath the Tully suggests that the contact of these units is an unconformity. Cross-section 2 extends southwest along the Lake Erie shoreline from the Eckhardt-Agle 1 well to Westfield, New York, northern Chautauqua County (Fig. 43). Noteworthy features in this section include: (1) the thickening of the Tichenor limestone to the west of Erie County; (2) the appearance of the Tully and underlying "Driller's Tully" in the Westfield area; and (3) the thinning of the interval between the base of the Tichenor limestone and the base of the Middlesex shale to the southwest (the Middlesex and Tichenor are separated by less than one meter of gray shale in the Muscarella I well; Fig. 43) followed by its rapid thickening as the Tully reappears near Westfield. This latter point is consistent with localized uplift. That the Middlesex shale and younger deposits show no evidence of this event provides clues to the timing of uplift, perhaps a reflection of localized tectonic activity of the post-Tully basin in advance of the Acadian thrust load and related forebulge (e.g., Tankard, 1986; Beaumont et al., 1988; Hamilton-Smith, 1993; Ver Straeten and Brett, 2000; Filer, 2003). Indeed, Filer (2003) defined a Famennian forebulge in West Virginia based on isopach and sedimentary facies maps that, when traced north, extends through the eastern Chautauqua County-Erie County region. We are just beginning to consider this event in a regional sense based on subsurface information from western New York state, northwest Pennsylvania, and northeast Ohio.
The Taghanic unconformity appears to be erosional throughout western New York state (Johnson, 1970; Wright, 1973; Baird and Brett, 2003). Heckel (1973) maintained that the contact is a slight diastem in central New York state, increasing in stratigraphic magnitude to the west where the disconformity represents an increasing period of non-deposition. However, our ongoing investigation of the subsurface stratigraphy of western New York state reveals a more complex situation made more so by the presence of the "Driller's Tully." Indeed, the initial event may have been that phase of erosion that occurred before deposition of the "Driller's Tully," which locally rests unconformably on the Tichenor limestone. Heckel (1973) pointed out that the Geneseo black shale conformably overlies the Tully limestone east of Cayuga Lake. West of this meridian, however, the contact becomes one of disconformity as progressively younger deposits onlap westward above the unconformity (de Witt et al., 1993) until the Geneseo shale feathers out in easternmost Chautauqua County into eastern Wyoming County (Fig. 42). Rickard (1989) suggested that the post-Tully unconformity was entirely subaqueous. Similarly, Baird and Brett (1991) postulated that the erosional contact of the Geneseo black shale and underlying Tully limestone is a submarine erosional feature (maximum flooding surface) related to the abrupt increase in water depth, likely the result of the sudden sinking of a forebulge in response to tectonic loading to the east (e.g., Hamilton-Smith, 1993).
III. Moving northeast along the beach for perhaps a kilometer enables us to observe in cliff exposure the entire Moscow shale - basal Rhinestreet shale interval (Fig. 40). Note the diagenetic carbonate intervals within the Cashaqua shale. The heavily fractured Rhinestreet black shale at the top of the exposure is easily recognized from the beach. The upper several meters of the Cashaqua and lower few meters of the Rhinestreet shale carry the older NS-trending joints, similar to what we saw on Pike Creek and likely for the same reasons. NW- and younger ENE-trending joints, most densely formed in the basal organic-rich interval of the Rhinestreet shale, formed as a consequence of catagenesis as these source rocks approached and finally reached their modeled maximum burial depth of ∼3.1 km (see Fig. 9). This exposure demonstrates the strong stratigraphic control on jointing in these rocks. The NW joints probably formed soon after the rocks had entered the oil window and at that time that the basin was feeling the same remote stress field responsible for emplacement of the Blue Ridge-Piedmont Megathrust. The younger ENE-trending joints propagated after the Rhinestreet shale had been in the oil window for a protracted period of time, thus their greater density. The stress field under which these joints formed reflected the oblique convergence of Laurentia and Africa and had been responsible for the generation of the early (perhaps Chesterian) ENE-trending coal cleat throughout the Appalachian Basin and similarly oriented joints in the deeply buried Frasnian black shales of the Finger Lakes District (Engelder and Whitaker, 2006).

STOP 2. EIGHTEENMILE CREEK/ROUTE 20 BRIDGE
Significance:
I) approximately the lower half of the Rhinestreet shale can be observed in the east wall of the creek exposure immediately south of the Route 20 bridge over Eighteenmile Creek;
II) the presence of NS joints at the contact of the Cashaqua gray shale and overlying Rhinestreet shale, and the high density of NW- and younger ENE-trending joints at the organic-rich base of the Rhinestreet shale;
III) basin modeling of measured vitrinite reflectance and the possibility of vitrinite suppression;
IV) recognition of the Cashaqua-Rhinestreet contact in the subsurface- gamma-ray vs. bulk density logs. I. A bit less than the lower half of the Rhinestreet shale is visible in the cliff exposure immediately north of the Route 20 bridge. We will have an opportunity to look at the concretion horizon - the middle concretion horizon- visible at the top of the cliff- ∼34 m above the bottom contact (see Fig. 1) - at Stop 4. The Rhinestreet section displayed in the cliff contains two intervals of gray shale and thin-bedded black shale at ∼10.0- 12.5 m and 27.5-29.3 m above the contact (Fig. 1). These horizons, which probably represent eustatic oscillations, can be observed in gamma-ray and bulk density logs (Fig. 44). Moreover, the lower gray shale interval contains a lenticular limestone bed as thick as 20 cm, one of Luther's (1903) scraggy layers and our lower scraggy layer (see Fig. 1 ). The geochemistry of the scraggy layers, discussed above, confirms their early diagenetic origin and suggests that they formed in a manner similar to the concretions. The scraggy layers, which are consistently found associated with organic-lean gray shale intervals (see Fig. 1), are recognized in gamma-ray logs as a sharp deflection to the left (Fig. 44) and may serve as markers for basinwide correlation. The fact that the limestone deflection on the gamma-ray logs cannot always be correlated among neighboring offset wells demonstrates the discontinuous nature of these deposits, which is evident in field exposures.
II. The Cashaqua shale-Rhinestreet shale contact is exposed beneath the bridge. The Cashaqua shale along Eighteenmile Creek comprises ∼11.5 m of silty, organic lean (TOC 0.7%) gray shale and nodular limestone horizons. The overlying Rhinestreet shale is composed of ∼54 m of organic-rich black shale, much lesser gray shale, sparse thin siltstone beds, and several carbonate concretion horizons and nodular limestone intervals, the latter termed "scraggy layers" by Luther (1903; see Fig. 1 for the measured Rhinestreet section along Eighteenmile Creek). TOC content of the Rhinestreet shale at its base is 8.09%, diminishing upward (Fig. 1; Lash and Blood, 2004b; Lash et al., 2004). Vitrinite reflectance of the Rhinestreet shale at its base along Eighteenmile Creek is 0.74%, indicating that these rocks entered the oil window during their burial history (Lash, 2006c).
Joints of the three dominant sets carried by the Rhinestreet shale are represented at this exposure. Several long NS-trending joints, can be seen at the contact. We contend that uplift of the Upper Devonian sedimentary succession in response to passage of a peripheral bulge at the onset of the Alleghanian orogeny, and resultant thermoelastic loading, induced the NS joints in the high modulus carbonate. These joints propagated upward under the influence of moderately overpressured pore fluid at the top of the Cashaqua shale. Elevated pore pressure may have resulted from a combination of disequilibrium compaction and the tight, impermeable nature of the organic-rich Rhinestreet shale. The lack of bioturbation in the finely laminated organic-rich sediment preserved a strong strength anisotropy that precluded easy vertical movement of fluid. The squeezing of ductile organic matter into voids and pore throats further reduced the permeability of the black shale (Lash, 2006b).
NW- and ENE-trendingjoints are densely formed in the black shale immediately above the contact with the Cashaqua shale (station EMC 13 and 16, Fig. 38). A similar relationship of joint density and TOC has been observed in the Dunkirk shale along the Lake Erie shoreline as well as in the Geneseo shale of the Finger Lakes District (Loewy, 1995; Lash et al., 2004). The correlation of joint density with TOC provides compelling evidence that the joints formed as natural hydraulic fractures as a consequence of the transformation of organic matter to hydrocarbons when these rocks were buried to the oil window during the Alleghanian orogeny (Lash et al., 2004; Lash and Engelder, 2005).
III. As noted earlier, we used the EASY%Ro kinetic model of vitrinite reflectance (Sweeney and Burnham, 1990) to model the burial/thermal history of the Rhinestreet black shale. Thermal modeling by use of the EASY%Ro algorithm requires knowledge of (1) the age(s) of the unit(s) of interest (the base of the Rhinestreet shale), (2) at least a partial thickness of the local stratigraphic sequence, and (3) the measured vitrinite reflectance of the unit(s) of interest (vitrinite reflectance of the base of the Rhinestreet shale = 0.74%). Our model for the Rhinestreet shale, which includes passage of the early Alleghanian peripheral bulge (see Fig. 9), requires ∼1.5 km of Permian strata in this area of the basin, an amount close to that modeled Beaumont et al. (1988) (Fig. 45).
Our method of modeling the burial/thermal history of the Rhinestreet shale depends on accurate %Ro values. However, several workers (e.g., Nuccio and Hatch, 1996; Rimmer and Cantrell, 1988; Rimmer et al., 1993; Rowan et al., 2004) have recognized the likely occurrence of vitrinite suppression in Devonian-Mississippian black shales, including those of the Appalachian Bain. Suppression of vitrinite leads to an underestimation of thermal maturity and, therefore, an underestimated burial depth. Suppression is most readily recognized by anomalously low %Ro on maturity profiles of wells or stratigraphic columns (e.g., Price and Barker, 1985). There are several lines of evidence suggesting that Devonian black shales of the Appalachian Basin have been affected to some extent by vitrinite suppression. Rimmer and Cantrell (1988), based on a vitrinite reflectance gradient derived from coal rank trends, claim that measured %Ro values of the Cleveland Member of the Ohio shale in eastern Kentucky are suppressed by as much as 0.3 - 0.5 %Ro. More recently, Rowan et al. (2004), in the course of their burial/thermal modeling of a cross-section from central West Virginia to northeast Ohio, demonstrated that the predicted %Ro of Devonian shale is locally twice the measured %Ro. Rowan et al. (2004) constrained their burial/thermal model with measured %Ro of Pennsylvanian coals and color alteration index values of conodonts recovered from Ordovician and Devonian limestones.
We suggest that Middle and Upper Devonian black shale of the western New York state and northwest Pennsylvania region of the Appalachian Basin was also been affected by suppression. Our data base includes published thermal/maturity values (Weary et al., 2000; Repetski et al., 2002) and%Ro and Rock-Eval data from our work on the Upper Devonian shale succession of the Lake Erie District. We limited our data to the western New York state and northwest Pennsylvania region of the basin in order to assume a common thermal history for these deposits. A plot of Rock-Eva! HI versus %Ro of Appalachian Basin black shale samples, principally from the Marcellus shale, reveals a trend similar to that observed by Nuccio and Hatch (1996) for the Upper Devonian-Mississippian New Albany shale, which has been affected by vitrinite suppression (Fig. 46). The higher HI samples, including samples collected from the Dunkirk and Rhinestreet shales, fall along a relatively narrow %Ro trend; however, at an HI of ∼125 mg/g, measured %Ro increases rapidly from ∼0.75 to 2.5 as HI drops to near zero (Fig. 46). This plot appears to show suppressed values in the lower to mid oil window(%Ro; = ∼0.7 to 1.0), consistent with Lo's (1993) findings regarding vitrinite suppression of hydrogen-rich organic matter. An especially interesting subpopulation of Marcellus data exists for the Finger Lakes District of New York state and adjacent Potter and Tioga counties, Pennsylvania (Fig. 46). These data are defined by %Roranging from 0.5 to 1.0 and HI 55 (Fig. 46). These samples, which appear severely suppressed, were collected from that part of the basin affected by emplacement of Cretaceous ultramafic intrusions (Kay et al., 1983).
Upper Devonian black shales exposed along the Lake Erie shoreline show a good amount of variation in measured %Ro over a relatively small distance (see Fig. 7). Measured %Ro in a well cutting sample collected from the base of the Rhinestreet shale in northern Chautauqua County is 0.51%. The %Ro of the same stratigraphic interval ∼40-45 km to the northeast is 0.73 - 0.77 based on the analysis of four outcrop samples. Finally, a sample collected from the base of the Rhinestreet 7.25 km to the east on Cazenovia Creek has a %Ro of 0.59. We recognize similar erratic variations in the measured %Ro for Dunkirk shale samples. Rowan et al. (2004) suggest that variations in the type of organic matter in marine sediments may account for erratic variations in %Ro (e.g., Ujiie et al., 2004).
Vitrinite suppression has also been attributed to the overpressuring of rocks (McTavish, 1978 1998; Carr, 2000). Pervasive natural hydraulic fractures in the Rhinestreet shale provides evidence that these rocks were overpressured during their burial history, which could have contributed to suppression of vitrinite reflectance. Regardless of its mode of origin, suppression of vitrinite reflectance is a phenomenon that should be taken into account in any exploration and production considerations of Devonian black shales.
IV. The contact of the Cashaqua gray shale and overlying Rhinestreet black shale is readily defined in the field; both color contrast and relative weathering profile provide easy means for making this recognition. The contact is rather obvious in gamma-ray logs, which provide a measure of the intensity of natural gamma-rays (uranium concentration) emitted by a formation (Fig. 44). Still, the initial deflection marking the contact is not what one might expect based on the abrupt contact between the low-TOC (<0.7%) Cashaqua shale and the carbonaceous (TOC>8%) Rhinestreet shale. However, the Cashaqua-Rhinestreet contact is very obvious in bulk density logs (Fig. 44), no doubt a reflection of the low density of organic matter relative to shale minerals (Smith and Young, 1964; Schmoker, 1979). Schmoker (1981) claimed that the bulk density log is a more accurate proxy for organic matter content than the gamma-ray API value. Indeed, the strong inverse relationship of bulk density and TOC in Upper Devonian black shale of western New York state has been confirmed by a proprietary data set we have had the opportunity to study (Lash, 2006c). It is interesting to note that the organic-rich Pipe Creek black shale (TOC = 3.2 - 7.9%) is not easily recognized on the gamma-ray log, but is readily picked on the bulk density log (Fig. 44). On the other hand, the base of the overlying Dunkirk shale is readily recognized on both types of logs (Fig. 44).
Luning and Kolonic (2003) have pointed out that many black shale deposits show a strong correlation of uranium, manifested by gamma-radiation in API units, and TOC. However, they also note those examples where the correlation is weak at best. The basal Rhinestreet shale and the Pipe Creek shale provide examples of the latter phenomenon. Indeed, a spectral gamma-ray log from a well in northeast Chautauqua County reveals the base of the Rhinestreet shale to be marked by an increase in uranium from 6 to 10 ppm; the Pipe Creek shale is defined by an increase of only 2 ppm. Luning and Kolonic (2003) suggest that the precipitation of uranium from seawater onto the surface of organic particles under anoxic conditions is governed by several factors, including carbonate content, sedimentation rate, and primary uranium content of the seawater. Upper Devonian shales of western New York state are defined by very low CaCO3 abundances ruling out the first factor. Luning and Kolonic (2003) point out that the longer organic matter is in contact with seawater, the greater the amount of uranium that can precipitate onto organic grains. Low sedimentation rates favor high uranium/TOC ratios. The anomalously low API values of the most organic-rich interval of the Rhinestreet shales (as deduced by TOC analysis of field samples and bulk density values determined from wireline logs) may indicate a high sedimentation rate of this interval such that lesser amounts of uranium were fixed onto organic particles. The same can be said of the Pipe Creek shale. The organic-rich base of the Dunkirk shale (Lash et al., 2004; Lash and Engelder, 2005), however, defined by the highest API value and lowest bulk density of the unit, would have accumulated relatively slowly.
Finally, it is possible that the introduction of low uranium freshwater into the Devonian Sea, such as might be expected as a consequence of the reduction of polar ice volume, was responsible for some of the variations in gamma-ray response observed in wireline logs of the base of the Rhinestreet shale and especially the Pipe Creek shale, which has been correlated with the Lower Kellwasser Bed recognized in Europe (Over, 1997). The concentration of uranium in seawater is an order of magnitude greater than freshwater (Broecker and Peng, 1982). The presence of the diagnostic uranium signature of the Pipe Creek shale in well logs recovered from northwest Pennsylvania, western New York state, northeast Ohio, and northernmost West Virginia suggests a basinwide cause. The Pipe Creek interval is a particularly interesting horizon. McGhee (1990) argued that this interval is a significant crisis interval similar to the younger Frasnian-Famennian boundary. In fact, he suggested that the Pipe Creek interval was more critical in terms of the maximum number of species extinctions and the extinction rate. It has been suggested that the Kellwasser Beds, including the Pipe Creek shale, resulted from eustatic transgressions perhaps caused by a reduction in polar ice volume (Sandberg et al., 1988; Buggisch, 1991 ; Joachimski and Buggisch, 1993). It may be that the same global mechanism that created the Pipe Creek crisis - rapid climatic warming - also accounts for the unique wire line log signature of the Pipe Creek shale.

STOP 3. STURGEON POINT
Significance:
I) comparison of the degree of jointing in black shale versus gray shale and the preferential generation of NS systematic joints in gray shale at contacts with overlying black shale;
II) carbonate concretions in black shale and analysis of compaction strain;
III) variation in the thickness of the bottom gray shale interval of the Rhinestreet shale.
I. Depart the parking area walking northeast toward the bedrock cliff a few hundred meters distant. We will follow this exposure of gray and black shale of the Rhinestreet shale for perhaps 0.75 km. The 2.75-m-thick gray shale interval exposed at the west end of the exposure is the equivalent of the lower gray shale interval of the Rhinestreet shale on Eighteenmile Creek (see Fig. 1; 9.9-12.5 m). The gray shale includes several horizons of high aspect ratio (lenticular) carbonate concretions and thin grayish-black shale layers. One can readily observe differences in the degree (density) of jointing between black and gray shale in this exposure. Black shale above and below the gray shale carries NW joints and younger ENE-trending systematic joints (Fig. 4 7). NW and ENE joints can be observed in the gray shale, but these deposits also contain the older NS trending systematic joints (Fig. 47). The black shale clearly is more densely jointed than the gray shale, a recurring theme throughout the Appalachian Plateau (Fig. 48; Loewy, 1995; Lash et al., 2004). Moreover, ENE joints are more densely formed (i.e., closer to saturation) than are the older NW joints. This latter observation can also be made of ENE and NW joints in the gray shale (Fig. 48). NS-trending systematic joints in the gray shale are more densely formed than are the younger NW joints. As described earlier, the NS joints are interpreted to have originated in the higher modulus carbonate concretions as a consequence of thermoelastic loading caused by passage of a peripheral bulge across the Appalachian Basin at the onset of the Alleghanian orogeny. Indeed, it is more common for a NS joint to penetrate or cut a higher modulus carbonate concretion than it is for ENE and NW joints to cleave concretions. Note also that very few NS joints penetrate more than a few tens of cm into the overlying black shale; more often than not, they are arrested at the base of the organic-rich deposits. This exposure presents a nice example of the stratigraphic control on jointing in these rocks.
Continue NE along the beach toward a large drainage pipe in the cliff. Careful observations of the gray shale horizon while working along the beach to this location would have revealed a gradual decrease in thickness of this interval. The thickness of the gray shale immediately east of the drainage pipe (determined by digging through beach sand to black shale) is ∼80 cm, markedly reduced from what is was at the start of this trek along the beach. Note also that the gray shale is not as complexly jointed here; indeed, the median spacing of NS joints is 50% greater than at the western end of the exposure (Fig. 48). Finally, the thickness of the gray shale interval at the NE end of the exposure is at least 1.5 m. NW- and NS trendingjoints are carried by the gray shale in this area. Further, NS joints are more densely formed at this location of the exposure than elsewhere along the beach (Fig. 48).
There is a school of thought that holds that all jointing in the exposed Devonian succession of the Appalachian Plateau of western New York state is a consequence of post-Alleghanian (perhaps even post-glacial) uplift of the Appalachian Basin. This view fails to take into account the well-defined chronology of the joints (Engelder and Geiser, 1980; Lash et al., 2004) and the fact that the joints provide information regarding the orientation of the remote or lithospheric (Engelder and Whitaker, 2006) stress field that was responsible for their formation. There are likely subtle differences in rock properties between the black and gray shale that perhaps played some ancillary role in the jointing of these rocks. We suggest that the black shale, principally by virtue of its greater amount of ductile organic matter, was not as stiff (i.e., lower modulus) as the gray shale, a view at odds with the more common thinking regarding heavily fractured black shale. Accordingly, gray shale, as a consequence of its higher porosity and locally greater abundance of calcite cement (e.g., Lash, 2006b) would have had the higher modulus. Joints driven by fluid decompression, natural hydraulic fractures, form at great burial depth and, therefore, under high confining pressure. All else being equal, it would have been more difficult to initiate open mode cracks in higher modulus organic-lean gray shale under these conditions. Moreover, movement of the black shale into the oil window and resultant transformation of kerogen to hydrocarbons would more readily create a negative effective stress necessary for the propagation of natural hydraulic fractures, their growth direction being controlled by the remote stress field. NW-trending joints were produced by an Alleghanian remote stress field that was responsible for emplacement of the Blue Ridge-Piedmont Megathrust sheet; ENE joints, however, reflect a longer-standing lithospheric scale stress field related to continued oblique approach of the African portion of Gondwana against the Laurentian plate boundary (Engelder and Whitaker, 2006).
The remaining question, then, relates to the preferential formation of NS joints in the gray shale as observed at this exposure. Passage of the peripheral bulge across the Appalachian Basin at the onset of the Alleghanian orogeny and related uplift of the Upper Devonian shale-dominated succession of the Lake Erie District subjected these rocks to thermoelastic contraction. The higher modulus rock - the gray shale would have carried a higher tensile stress and, therefore, would have been closer to the negative effective stress necessary to initiate joints. The inferred contrast in modulus between black and gray shale may explain why the NS joints terminate at gray shale-black shale contacts; i.e., the lower modulus black shale carried a lower tensile stress than the gray shale. However, the highest modulus rocks in the Rhinestreet shale were the carbonate concretions. As described earlier, concretions remained in compression during catagenesis so that fluid-driven NW and ENE joints were unable to penetrate them any more than a cm or so. However, during thermoelastic contraction, the concretions carried a high level of tensile stress (higher than that of the gray shale and certainly much higher than that of the black shale) making them more likely to fail under true tension. Initiation of the NS joints in the diagenetic carbonate was followed by their propagation through the host gray shale aided by a modest pore pressure.
In sum, this exposure, like others in this region of the Lake Erie District, provide evidence that systematic joints carried by the Upper Devonian shale succession resulted from two different loading mechanisms before post-Alleghanian uplift of the Appalachian foreland. However, we are investigating the intriguing possibility that an ENE-trending set of joints, some of which cut diagenetic carbonate, formed as a consequence of the uplift of the Appalachian Plateau. These joints appear to have been driven by thermoelastic contraction under a remote stress field perhaps induced by spreading at the Mid-Atlantic Ridge; thus their ENE strike.
II. This exposure also allows for the differentiation of concretionary carbonate found in gray shale from concretions common to black shale (spectacular examples of this latter type will be seen at Stop 4). The former are defined by high aspect ratios (they may better be described as lenticular diagenetic limestone) and lack obvious internal structures, principally layering. The scraggy layers are an example of this type of diagenetic carbonate. Carbonate concretions in the black shale have lower aspect ratios and display nicely preserved layers of the host organic-rich sediment. The above distinctions is readily made at this exposure.
Numerous small, internally laminated concretions can be observed in black shale within a meter or so of the upper contact of the gray shale interval, especially along the eastern stretch of the exposure. These concretions lend themselves to compaction strain analysis as described earlier (Fig. 49). Compaction strain is determined by measuring a stratigraphic interval within the concretion, Ti, and that same interval traced into the host shale but far enough beyond the strain shadow to provide a full measure of compaction strain, To (Fig. 49). Estimated compaction strain of six concretions at this exposure is 60.8%, ∼9% higher than the average for the Rhinestreet shale as discussed earlier. We believe that this reflects concretionary growth in a stratigraphic horizon that had not compacted to the extent that most of the unit did before the onset of normal compaction (see earlier discussion).
III. Moving SW to NE along the beach, we observed the thickness of the gray shale interval diminish from ∼2.75 m to ∼0.8 m; farther along, though the thickness increases to at least 1.5 m. We observed no evidence of faulting that might account for the described change in thickness. Perhaps we are looking at the effects of differential compaction (e.g., Conybeare, 1967; Finely, 1990; Ryder, 1996). Indeed, the local presence of flame-like structures at the top of the gray shale interval suggests that these deposits were rather mobile and perhaps characterized by moderately high pore pressures, as suggested by the NS joints. Indeed, the relatively low density of NS systematic joints in the middle interval of this exposure where the gray shale is thinnest may reflect the effects of differential dewatering and resultant differential compaction. Preferential dewatering of the gray shale could account for (1) the reduced thickness of the gray shale at this location of the exposure and (2) a reduction in pore pressure, which would not have favored the generation of NS-trending joints in these rocks. Clearly, much more work is needed to elucidate this hypothesis.

STOP4. EIGHTEEN MILE CREEK (main branch)
I) the contact of the Rhinestreet shale and overlying Angola gray shale;
II) general characteristics of the carbonate concretions.
I. Follow the creek upstream (to the left) to where the east branch of Eighteenmile Creek enters the main creek. Cross the east branch just upstream from this point and follow a trail through the woods for roughly several hundred meters or until the path approaches the main branch. Cross the creek and walk to where you can view the cliff exposure across the creek. The middle concretion horizon will be visible a meter or so above the waterline. The upper ∼21 m of the Rhinestreet shale and lower few meters of the Angola shale are exposed in the cliff (see Fig. 1). The Rhinestreet-Angola contact is transitional and marked by an increase in the frequency of less resistant (to the effects of weathering) gray shale layers, which can be observed both in the field at this stop and in well logs (Figs. 1 and 50). Luther (1903) suggested that the contact be placed at an easily recognizable discontinuous nodular limestone bed locally as thick as ∼30 cm, the upper scraggy layer (see Fig. 1 ), which can be seen near the top of the exposure. The upper scraggy layer can also be recognized in some gamma-ray logs by a sharp deflection to the left (e.g., Fig. 20).
II. Carbonate concretionary horizons like the one exposed here are conspicuous in Middle and Upper Devonian black shale units of central and western New York state (Dix and Mullins, 1987; Lash and Blood, 2004a,b), southwestern Ontario (Daly, 1900; Coniglio and Cameron, 1990) and Ohio (Clifton, 1957; Criss et al., 1988) suggesting that diagenetic conditions favorable to the formation of these features existed regionally. Concretions are especially useful in the analysis of the early post-depositional history of encapsulating shale, for not only can they preserve the depositional texture of the host sediment (e.g., Woodland, 1984; Duck, 1990), but they carry a record of the evolution of pore fluids carried by the sediment (e.g., Hudson, 1978; Coniglio and Cameron, 1990). Details of our work on the carbonate concretions of the Rhinestreet shale can be found in Lash and Blood (2004a,b) and are presented in an abbreviated format above.
The concretions visible here are among the most spectacular in the world. They are similar to those described from such well known units as the Kimmeridge Clay (Astin and Scotchman, 1988) and the Jet Rock (Hallam, 1962). Close inspection of the exposed Rhinestreet concretions reveals their laminated nature and presence of septaria. Our principal interest in the concretions has been their use as compaction strain indicators. However, before they could be used in this way, we had to understand their mode of origin and depth of growth, which was considered earlier. Most carbonate concretions of the Rhinestreet shale are found in three stratigraphically confined but laterally persistent horizons (see Fig. 1) - a lower horizon- 6 m above the base of the unit, a middle interval slightly more than halfway up the Rhinestreet (Stop 4), and an upper, less well defined, horizon ∼43 m above the base of the unit. Most concretions are oblate ellipsoids with maximum diameters and thicknesses ranging up to 2.7 m and 1.1 m, respectively (Lash and Blood, 2004b); coalesced concretions have also been observed and are locally abundant. We carried out a detailed analysis of the dimensions and general shapes of concretions of the lower and middle horizons (Fig. 51). The dimensional data for both horizons reveals an especially intriguing feature; the maximum vertical dimension of the bulk of the concretions is ∼80 cm (Fig. 51). However, the scatter of data from the middle horizon suggests that some concretions stopped growing in the vertical dimension at ∼80 cm but continued to form laterally (Fig. 51). This suggests that the middle concretion horizon remained in the zone of anaerobic methane oxidation (AMO) longer than the lower concretion horizon.
Crucial to precipitation of diagenetic carbonate by the AMO mechanism discussed earlier is a marked reduction in sedimentation/burial, which would stabilize the sediment in the AMO zone thereby enabling protracted carbonate precipitation in that horizon (Raiswell, 1987, 1988). Two observations considered earlier provide evidence of a marked reduction in burial rate during concretion growth: (1) little or no change in carbonate volume percentage outward from the centers of studied concretions and (2) negligible variations in laminae thickness across concretions (Lash and Blood, 2004b).
The middle and lower concretion horizons differ in a fundamental way that yields indirect information regarding such parameters as the magnitude and duration of the change in sedimentation/burial rate. The middle horizon, 1 to 2.5 m-thick, contains concretions distinctly larger in the horizontal dimension than are concretions of the roughly 2.5-m-thick lower horizon (Fig. 5 1). The thickness of a concretion horizon is interpreted to be proportional to the change in sedimentation/burial rate; i.e., a marked reduction in burial rate yields the minimum vertical range of carbonate precipitation because the zone of AMO in the sediment stabilized (Raiswell, 1987, 1988). The more oblate, occasionally flat-topped, concretions in the middle horizon, then, likely reflect the restriction of carbonate precipitation to a relatively narrow, fixed zone defined by the upper limit of methane and CO32- diffusion and the maximum depth to which seawater sulfate can diffuse (Raiswell, 1988). Thus, the generally larger horizontal dimensions of concretions of the middle horizon indicate that the break in sedimentation persisted longer at this time than during formation of the lower concretionary horizon.
The middle Rhinestreet concretion horizon, by virtue of its limited thickness and the geometry of its concretions, records a reduction in sedimentation/burial rate of greater magnitude than that which led to formation of the thicker lower concretionary interval (Lash and Blood, 2004b). However, the thickness and geometry of the scraggy layers likely reflects an even greater duration of the break in sedimentation/burial and, therefore, a protracted period of time that isotopically light methane and 13C-enriched dissolved carbonate were delivered to the stationary AMO zone (Raiswell, 1987, 1988). That is, the near cessation of sedimentation and, more importantly, subsidence, resulted in confinement of carbonate precipitation to a narrow, fixed zone defined by the upper limit of methane diffusion and the maximum depth to which seawater sulfate could diffuse (Raiswell, 1988). Because the events that contributed to the reduction in sedimentation/burial rate may have been basinwide, concretionary horizons may serve as time lines in establishing regional time-stratigraphic relationships.

B2: SILURIAN SEQUENCE STRATIGRAPHY, EVENTS, AND PALEOENVIRONMENTS ALONG THE CRATONIC MARGIN OF THE APPALACHIAN FORELAND
Carlton E. Brett
University of Cincinnati
Cincinnati, OH 45221-0013
(carlton.brett@uc.edu)
Patrick I. Mclaughlin
Geology Department
Bucknell University
Lewisburg, PA 17837
R.Jacobi
876 NSC
Silurian rocks of the Appalachian Basin and the North American mid-continent platform provide an excellent suite of strata for application of sequence, and event stratigraphic approaches. The strata are well exposed and display marked vertical changes in facies, commonly associated with distinctive condensed beds and/or discontinuities. The Niagara Escarpment in western New York and the Niagara Peninsula of Ontario represents a key reference area for the North American Silurian System. Indeed, the term "Niagaran," which has been variably applied to the middle or lower and middle portions of the Silurian, is still commonly used in North America. These rocks have been well documented by many researchers, beginning in the mid 1800s (Hall, 1852; Grabau, 1901; Foerste, 1906, 1935; Williams, 1919; Gillette, 1947; Bolton, 1957; Zenger, 1965, 1971; Sanford, 1969; Martini, 1971; Rickard, 1975; Duke, 1987; Duke and Brusse, 1987; Duke and Fawsett, 1987; Brett, 1983 a, b, 1985; Brett et al., 1990, 1995, 1998; Goodman and Brett, 1994; LoDuca and Brett, 1994).
Previous reports on the Silurian in the Niagara region have focused on identification of depositional sequences and depositional/faunal events. A theme of this guide is the documentation and interpretation of Silurian facies, events, and sequences along the northwestern rim of the Appalachian Basin. As such, the discussion is not restricted to the sequences and facies of western New York. A remaining research agenda is to trace these sequences into other regions, especially the mid-continent. Ongoing research along the Cincinnati Arch, in Ohio, Indiana, and Kentucky has revealed striking similarities of sequences and events among these seemingly disparate areas, some 500 to 700 km southwest of the main Niagara Escarpment (Fig. 1). In this paper we review and update information on depositional sequences and events in western New York and Ontario and then discuss their lateral relationships with those of the Cincinnati Arch. A companion field guide will highlight the Silurian strata of the Cincinnati Arch region. The large-scale ("third order") sequences described herein are sharply bounded stratal packages, ranging from less than a meter (where partially truncated or condensed) to about 50 m in thickness. In each case, the sharp basal contacts, when traced in an up-ramp direction, change from facies dislocations (abrupt shallow over deeper facies change) to erosional surfaces that demonstrably truncate underlying strata. Most sequences recognized herein, display a generally deepening-shallowing pattern.
However, in detail, the larger-scale sequences are composites, divisible into smaller, fourth-order sequences.
With its renewed emphasis on through-going discontinuities and condensed beds, the sequence stratigraphic approach, has encouraged a broader, more regional view of stratigraphy, and an attempt to understand the genetic significance of particular beds and surfaces. To some degree it vindicates the earlier "layer cake stratigraphy" approach. Sequence stratigraphy, originally developed from remote seismic studies of passive margin sediment wedges (Vail et al., 1977, 1991; Wilgus et al., 1988; Coe, 2003; Catuneanu, 2002), is now being applied at the outcrop scale to diverse depositional settings such as the Taconic foreland basin (Brett et al., 1990 a, b; Witzke et al., 1996; Brett et al., 1998; Brett and Ray, 2006). Many distinct horizons in the local stratigraphic record are interpretable as sequence boundaries, drowning surfaces, and forced regression surfaces. Moreover, a number of phenomena, which occur non-randomly in the geologic record, from phosphatic nodule horizons to reefs, fit in predictable ways into depositional sequences.
A second theme of this article is the recognition of widespread events, such as storm deposits, (tempestites), rapidly buried fossil horizons (obrution deposits), and even seismically deformed beds (seismites). The Silurian Period has typically been considered a tectonically quiescent time; however, growing evidence from eastern North America suggests the Silurian was more dynamic then previously considered. For example, sedimentology and stratigraphic mapping of Silurian deformed beds in eastern North America demonstrate that these event beds are extremely widespread and that their component sediment layers were not deformed during initial deposition, but slightly later, during shallow burial, likely the result of shaking induced by large-scale earthquakes (McLaughlin and Brett, 2006). The distribution of deformed beds, together with increased subsidence, clastic influx, and K-bentonite horizons, provides a meter of intensity and timing of pulses of Silurian orogenesis. A variety of such catastrophic events is recorded in the Silurian rocks of the Niagara region and will be examined on this trip.

B2: SILURIAN SEQUENCE STRATIGRAPHY, EVENTS, AND PALEOENVIRONMENTS ALONG THE CRATONIC MARGIN OF THE APPALACHIAN FORELAND

Northern edge of ridges of Barre Moraine
Northern edge of ridges of Barre Moraine and possibly the Albion Moraine, which converges with it here.

Cross Erie/NYS Barge Canal.

STOP A. (OPTIONAL) Crain Street Quarry and Outwater Park
(brief stop).
View off Niagara Escarpment; spillways (occupied by East and West Jackson Roads) of ancient Lake Tonawanda into preglacial Lake Iroquois. Note Lake Ontario (about 18 km to north).
Walk back from overlook into old quarry at brink of escarpment to observed glacially polished surface of Pekin Member, Gasport Formation with small bioherms with well preserved stromatoporoids, favositids, and colonial rugose corals; pelmatozoan holdfasts are attached directly to bioherm surfaces. Patches of cross-bedded reef-flank debris are present at margins of bioherms.

STOP 1. Road Cut on Gooding Street
Below Somerset Railroad Viaduct (Queenston Shale-Medina Group)
This excellent outcrop has been described previously in considerable detail (see Friedman, 1982; Duke et al., 1987). It provides an outstanding exposure of the basal Silurian Cherokee Unconformity and a good opportunity to study the lower units of the Medina Group as well as the uppermost beds of the Queenston Shale.
About 6 m (20') of upper red mudstones and siltstones of the Queenston Formation (upper Ordovician. Ashgillian) are exposed at this locality. The Queenston has been interpreted as either very shallow marginal marine or non-marine red beds. The Cherokee unconformity in this area is the basal surface of Whirlpool Sandstone, which is nearly planar. It is also a megasequence boundary separating the early or Creek phase of Sloss' (1963) Tippecanoe megasequence from the later Tutelo phase see Dennison, (1989).
The Medina Group (Sequence I of the Silurian System), consists of an Early Silurian (early Llandovery, Rhuddanian), siliciclastic wedge derived from tectonic source areas to the southeast (Figs. 5, 21). The lowest Silurian unit, white Whirlpool sandstone is about 3.5 m (11.5') thick at this location. Basal beds of the Whirlpool Sandstone are quartz arenites with northwest dipping cross strata, which have been interpreted as non-marine, braided stream deposits (Middleton et al., 1987). Large-scale channel-like structures occur in, or at least at the top of, these sands. Shale drapes within such channels at Lockport have yielded marine acritarchs (M. Miller, unpublished data) indicating that the channels were backfilled by very shallow marine sands and minor muds during a rise of sea level. Hence the irregular channeled surface that separates lower Whirlpool braided fluvial facies from upper Whirlpool hummocky cross-stratified, sparsely fossiliferous beds is a transgress surface. The Whirlpool Sandstone thus is interpreted to contain either a lowstand (or shelf margin) systems tract and a transgressive deposit. A thin bed containing phosphatic pebbles and fossil grains occur with the Whirlpool-Power Glen contact. This phosphatic pebble bed may mark a marine flooding surface, or surface of maximum starvation associated with relatively increased rates of sea level rise. This surface marks the change from shallow shelf sands of the upper Whirlpool into deeper shelf muds and storm sands of the upper Whirlpool into deeper shelf muds and storm sands of the Power Glen Formation, herein interpreted as highstand deposits. The Power Glen exhibits subtle small-scale parasequences.
This outcrop is one of the easternmost exposures of the Power Glen Shale. At this locality the Power Glen Shale comprises of about 5 m (16') of greenish gray shale with thin tempestitic siltstone and sandstone beds. The basal meter-thick transitional zone consists of thin (2-10 cm) muddy sandstones with interbedded sandy shales. Sandstone in the Power Glen Shale feature small-scale hummocky lamination and gutter casts suggestive of shallow, storm influenced shelf deposition. Small burrows (Planolites) are common, but body fossils are rare.
Greenish to reddish sandy shales and reddish sandstones occur near the top of the Power Glen suggesting a minor upward shallowing trend. However, the top of the unit (as defined herein) is sharply demarcated at the base of a massive white to pink mottled sublitharenitic sandstone about 2.5 meters (7.7') thick, the Devil's Hole Sandstone (Brett et al. 1995) seen at Niagara Gorge (Stop 6). The basal and upper beds of the sandstone contain lingulid brachiopods and probable Lingula burrows.
The white sandstone appears to record relative sea-level drops during which sands were distributed widely into the basin. The unit has some characteristics in common with the upper member of the Whirlpool Sandstone and, by analogy, is considered to be a relative lowstand to transgressive deposit.
The unnamed sandstone, in turn, is overlain by about 2.0 meters of brick red shales and interbedded sandstones assignable to the lower Grimsby Formation. These beds are ferruginous and exceedingly rich in fragments of lingulids with rare nautiloids and bryozoans. Thin spastolithic (oolitic) hematite stringers occur near the top of the unnamed sandstone and probably reflect reworking of sediments in shallow marine environments during an interval of sediment starvation. Hence, these shell-rich ferruginous sediments represent a condensed interval at the base of the Grimsby highstand deposits.
The reddish marine shales near the base of the Grimsby Formation pass upward into red and white-mottled sandstones and thin sandy shales. These beds are exposed high in the cut and are not readily accessible. This upper interval will be seen to better advantage at Niagara Gorge (Stops 4,6).

STOP 2A. Lockport Junction Road Cut (Lower)
Upper Medina and Lower Clinton Groups
This cut along Lockport Junction Road exposes upper units of the Medina Group and the lower part of the Clinton Group. The lowest beds are exposed beneath and just north of the overpass of Lower Mountain Road over Route 93. The basal units seen here are red shales near the top of the Grimsby Formation. These shales are overlain by 1.0 to 1.2 meter thick, blocky, pinkish gray sandstone that displays color mottling due to bioturbation. Swirly spreiten of the trace fossil Daedalus occur sporadically near the top of the sandstone ledge. Detailed regional correlation by Duke and Fawcett (1987) indicates that this unit is the equivalent of the Thorold Sandstone at the Niagara Gorge. The upper contact of the sandstone is marked by a thin red silty bed containing a "hash" of lingulid shell fragments.
The Thorold, in turn, is overlain by 1.7 to 1.8 m of dominantly red silty shale of the (Cambria Member) uppermost Medina Group. This shale bears distinctive fine, white mottling due to bioturbation. Although previously assigned to the Grimsby Formation, this is a distinctive, ostracode-bearing shaly unit that overlies the Thorold Sandstone and is traceable regionally at least as far as east as the Rochester area. Brett et al. (1995) termed this unit Cambria Shale Member of the Thorold Formation, using this section as the type locality. The upper portion of the Cambria Member is pale purplish to greenish gray sandstone and sandy shale that was formerly termed Thorold Sandstone. In actually, this is simply a leached sandy zone in the Cambria Shale. The Kodak Sandstone and about 1 m of Cambria Shale have been removed here at the sequence I/II unconformity. Sandstones contain small Skolithos burrows and intercalated green shale beds, especially the topmost layer, contain prolific leperditiid ostracodes. The greenish color extends down about 30 to 50 centimeters below the upper contact where sandstones are mottled pale purple and green.
This discoloration is probably associated with the top unconformity and deposition of overlying reducing sediments (Duke and Fawsett, 1987).
Here the top of the Medina Sandstone is an erosion surface overlain by a thin (3 to 5 centimeter) dark gray, phosphatic sandy dolostone (Densmore Creek Bed) with prolific Hyattadina brachiopod valves. A thin laminated siltstone rests on the bed at the contact with the greenish gray Neahga Shale, which at this locality is about 1.0 meter thick.
At the base of the Hickory Corners Limestone (Reynales) is a thin (3 to 5 cm) pyritic, sandy limestone packed with black phosphatic pebbles and shell fragments (Budd Road phosphatic Bed). It is overlain by about 40 cm of alternating greenish gray shales and thin limestones capped by a 60 cm-thick ledge of nodular crinoidal pack- and wackestone at the top of the road cut. These beds contain a prolific fauna including corals, brachiopods (Hyattidina, Dalejina, Platystrophia) and pentameric crinoid stems belonging to a newly described disparid, Haptocrinus. The upper surface at this locality is a glacially striated pavement and the post-Reynales erosion surface cannot be observed.

STOP 2B. Lockport Junction Roadcut (Upper)
(Upper Rochester, Decew, Gasport Formations)
This large road cut displays the upper part of the Rochester Formation (about 5 meters), DeCew Dolostone (2.5 meters), and Gasport Limestone (over 7 meters); it was described in details as Stop 1 of Brett (1982 NYSGA Guidebook). However since that time the exposure has been freshly blasted to widen Route 93.
A subsequence boundary occurs between the dolomitic calcisiltites and shales of the upper Rochester Shale and the overlying DeCew Dolostone. The latter is a buff-weathering, silty dolostone with thin layers of intraclasts and convoluted bedding. The basal contact is sharp and locally channeled, but nearly conformable. The upper Rochester and DeCew are burrowed in some levels, body fossils are rare but this exposure has produced long columns of crinoids in the lower DeCew.
The sharp and undulatory upper contact of the DeCew Dolostone with the Gasport Formation forms the boundary between the Clinton and Lockport groups, and is interpreted as a sequence bounding unconformity (between sequences V and VI). The surface represents an abrupt lowering of relative sea-level and beveling of older strata. The basal Gasport is a greenish gray brachiopod-rich, crinoidal grainstone conglomerate with dolostone clasts eroded from the DeCew. Missing at this contact is the newly recognized Glenmark Shale, a fossiliferous gray shale with litho and biofacies resembling the Rochester Shale.
Only the lower (Gothic Hill) member of the Gasport is present here. At this locality and at cuts along adjacent Gothic Hill Road this unit is exceptionally thick (7-8 meters) and composed of well to poorly sorted pelmatozoan pack- and grainstones. This facies is interpreted as a high-energy crinoidal bank. Brachiopods are common in an argillaceous, thin-bedded unit near the top of the unit. A small bioherm composed of algal and bryozoan boundstone (micrite) occurs within the Gothic Hill Member. The overlying units will be examined at the next stop.

Junction Thrall Road on right.
Exposures of fossiliferous Rochester Shale along Thrall Road were described as Stop 2 in Brett (1982).

STOP 3. "PEKIN BIOHERM" CUT ON RT. 429
Location: Cuts on both sides of NY 429, just north of underpass beneath Upper Mountain Road, Pekin, Niagara County, NY (Cambria 7.5' Quadrangle).
Description: This classic cut in the brow of Niagara Escarpment has recently been widened, and provides an excellent, but very fresh exposure of the upper Gasport (Pekin Member of Brett et al., 1995) and the lower portion of the overlying Goat Island Formation (Fig, 19). The large, massive mound exposed particularly on the west side of Rt. 429 was originally described as a Gasport bioherm and was considered to show reefal succession (Crowley and Poore, 1974). However, recent excavations have revealed that the mass of stromatoporoid bearing rock is almost entirely within the Goat Island Formation and is separated from older Gasport deposits (including a formerly exposed small bioherm, now destroyed by blasting) by a major erosion surface. The contact between the gray, thin bedded argillaceous dolostones of the Pekin Member of the Gasport and overlying massive, dolomitic crinoidal grainstone of the Goat Island (Niagara Falls Member of Brett et al., 1995), a sequence boundary, appears sharp, but horizontal, in the northeastern portion of the cut. However, this contact abruptly descends to near road level just north of the Upper Mountain Road overpass. Here a mass of biohermal lithology overlies the unconformity. On the freshly blasted west side of the cut the irregular contact between the dark gray Pekin and light pinkish biohermal Goat Island is now very clear. The erosion surface separating the units has a relief of over 2 meters. The Goat Island "bioherm" appears to have developed in low areas on the unconformity. It consists of a mixture of crinoidal grainstones and light gray dolomicrite, rich in stromatoporoids, many of which are tumbled onto their sides or even inverted, indicating storm disturbance. This biohermal mass built up in "channels" along a sequence boundary during initial transgression.

Good view of Niagara Gorge to right

Pass forebay on left; driving over sluiceways of Robert Moses Power Plant

Pass Niagara University on left

Pass by fisherman's parking area
Pass by fisherman's parking area (cars can be left here; we will continue on South Haul Road to the bottom of the gorge and allow passengers to disembark and return uphill on foot); ahead through underpass beneath Robert Moses Parkway are excellent views of the Niagara Gorge and a complete stratigraphic section from the mid Silurian Lockport Group down to the unconformable contact between the Silurian Whirlpool Sandstone and the underlying Ordovician Queenston Shale.

Fishermen's access at bottom of Niagara Gorge
Fishermen's access at bottom of Niagara Gorge, just outside gates for Robert Moses Power Plant; passengers will disembark and proceed back up the South Haul Road; vehicles will return back up the road to park

STOP 4. Niagara Gorge - South Haul Road
Location: Large road cuts in east wall of Niagara Gorge along South Haul access road for Robert Moses Power Plant, and ascending for about 1 km south to a tunnel beneath Robert Moses Parkway (Fig. 21). Parking is available in a fisherman's access parking lot just west of Rt. 62 (Hyde Park Boulevard) immediately south (uphill) from the tunnel. Lewiston, Niagara County, New York (USGS Lewiston 7.5' Quadrangle). Note: Access to the Haul Road exposure is strictly controlled by the Robert Moses Power Project and requires advance permission.
Description: This outstanding outcrop of Lower to mid Silurian strata shows important contrasts with sections of the comparable interval near Hamilton. The section begins near the Power Plant with about 8 m of the Upper Ordovician Queenston Shale (Fig. 21). Its sharp contact with the overlying Whirlpool Sandstone is the Cherokee Unconformity (Silurian Sequence I boundary).
The units of the Silurian succession are described in ascending order, as follows:
Medina Group (Fig. 21).
Whirlpool Sandstone: (4.5 m)- White, trough cross bedded, quartz arenite facies which record a non-marine to marine transition. Excellent profiles of channels are visible.
Power Glen Shale: (∼8 m)- Dark gray, friable shale, with very minor sandstone interbeds
Devils Hole Sandstone: (2m)- Pale gray, massive, quartz arenite with a distinctive, meter-thick phosphatic, sandy dolostone, Artpark Phosphate Bed near the top (IC).
Grimsby Formation: (15 m)- Greenish gray to maroon shales and mudstones with bundles of thin reddish and white mottled sandstones (Fig. 21).
Thorold Sandstone: (2m)- White, cross-bedded quartz arenite. The Thorold has a sharp, erosive base which marks the base of the next Medina subsequence (IC). A thin (2-10 cm), sandy phosphatic bed, the Densmore Creek Bed (Brett et al., 1995) rests sharply on the Thorold (and on a Cambria Shale remnant north of the Power Plant), marking the base of the Neahga Shale. Clinton Group (Fig. 18)
Neahga Shale: (2m)- Dark greenish gray, very friable shale (base sharp (II SB). Reynales Formation (Hickory Corners Member): (∼50 cm)- Medium gray, nodular, burrowed, bryozoan-rich wacke- to packstone. Conodonts indicate a mid Llandovery age for the Reynales (see LoDuca and Brett, 1994); this unit represents an erosional remnant of the Reynales.
Rockway Formation: (3 m)- Buff-weathering, argillaceous dolostone with thin dolomitic shales shows prominent rhythmic bands (10-50 cm) of sparsely fossiliferous argillaceous dolostone interbedded with thin gray shales. The Rockway shows a sharp upper contact (S-IV-V SB).
Irondequoit Formation: (2.5 m)- Massive, pinkish-gray, crinoidal pack-and grainstone. Clasts of fine-grained dolostone, derived from the underlying Rockway occur in the basal thin bed of the Irondequoit (Fig. 23). Its sharp upper contact (MFS) is marked by a 30 cm thick shell bed.
Rochester Shale: (18 m)- Medium dark gray mudstone with thin calcisiltites and lenticular fossil rich limestones; 1.5 m of bryozoan-rich limestone beds underlie the sharp top (MFS) of the Lewiston Member. The upper Rochester (Burleigh Hill Member) also displays a sharp contact (SSB) with the enterolithic DeCew Dolostone beds.
DeCew Formation: (3 m) - Dark gray, buff weathering, laminated dolostone (calcisiltite). Here and, especially in weathered exposures in the adjacent Devils Hole Park section, the DeCew displays spectacular soft- sediment deformation with isoclinally folded beds (seismite?).
Lockport Group (Fig. 24): The sequence V-VI boundary at the DeCew-Gasport contact is well exposed near the entrance to the "tunnel" beneath the Robert Moses Parkway at the top of the cut; basal Gasport shows rip-up clasts of DeCew Dolostone (Fig. 17).
Gasport Formation: (5 m)-Pinkish gray thin bedded to massive dolostone, divided into a lower pinkish gray dolomitic crinoidal grainstone (Gothic Hill Member) and a 2.5 m upper argillaceous, bioturbated dolostone (Pekin Member) Weathered surfaces of the Gothic Hill grainstones display probable bipolar cross- stratification . The sharp upper contact is a subsequence boundary.
Goat Island Formation: (∼10 m) - Buff weathering dolomitic, crinoidal grainstones, buff, thin bedded dolostone with white chert, and dark brownish gray, argillaceous, banded dolostone; the basal unit (Niagara Falls Member) is massive crinoidal grainstone with scattered Cladopora corals and stromatoporoids; abundant vugs appear to be solution cavities in a stromatoporoid-rich zone. The Lancaster Member is poorly developed here, thin (∼2.5 m) and only sparingly cherty. Argillaceous and bituminous gray dolostones of the Vinemount Member form the uppermost unit on the access road. This member is much less shaly than at its type area near Hamilton.

STOP 5. Robert Moses Access Road and Forebay
Location: Access road to the Robert Moses Power Plant, and adjacent forebay canal just west of Military Road and 1.2 km north of Route 31, Lewiston, Niagara Co., NY (Lewiston 7.5' Quadrangle).
Description: Higher units of the Lockport Group are visible along the forebay and consist of the upper Eramosa (new usage in New York) and lower Guelph formations (Fig. 24). Large algal bioherms characterize the uppermost units of the Eramosa Formation in the forebay. Exposures in the small road cut at the underpass of the access road beneath the lanes of I-190 show exceptionally large (2 m high) stromatolites, and some non-laminated thrombolites. The stromatolitic interval at the base of the Guelph is traceable with subsurface data at least to Hamilton (Fig. 24).

Crossing east end of Robert Moses forebay area
Pump Generating Power Plant to right; note small thrombolitic mounds in Eramosa Formation exposed in forebay cuts to left

Brow of Niagara Escarpment
On clear days good views of the mouth of Niagara River and Lake Ontario to the north; cuts to right in Gasport-Goat Island formations

STOP 6. Niagara Gorge, Lewiston (Artpark)
Location: Sections in east wall of Niagara Gorge along old haulage road extending from north end of Niagara Gorge at Niagara Escarpment just south of Artpark, off Fourth Street, southward to the Lewiston-Queenston bridge, Lewiston, Niagara County, NY (Lewiston 7.5' Quadrangle)
Description: The north-facing cuesta (Niagara Escarpment) stands 76 m above the adjacent Lake Ontario plain. Niagara Falls was initiated here about 12,000 BP.
Exposures of the Upper Ordovician Queenston Shale and its unconformable contact with basal Silurian Whirlpool Sandstone are visible along a short path, adjacent to the river, immediately south of the Artpark theater. Outcrops of the Lower Silurian stratigraphic units above the Whirlpool Sandstone are accessible along an old haulage road that leads southward from the Artpark Visitor Center into the gorge. At the entrance to the gorge (edge of Niagara Escarpment) an isolated "butte" of Lower Silurian strata (Power Glen-lower Grimsby; type locality of the Artpark Phosphate bed) between the path and the river, represents a remnant of a promontory in the gorge wall that was breached during excavation for the power plant. A 0.5 m layer of mottled sandstone with prominent ball-and-pillow structures, about 5.5 m above the base of the Grimsby, is well exposed in the cliffs about 200 to 700 m north of the Lewiston-Queenston Bridge. Caution is required in this exposure, as rock falls are common. Abundant fallen debris also provides an excellent look at varied lithologies of the upper Medina, as well as Clinton and Lockport units.

B3: AN EARLY LATE DEVONIAN BONE BED-PELAGIC LIMESTONE SUCCESSION: THE NORTH EVANS-GENUNDEWA LIMESTONE STORY
In memory of Daniel B. Sass (1919-2006)
Gordon C. Baird
Dept of Geosciences
SUNY-Fredonia
Fredonia, NY 14063
William T. Kirchgasser
Dept. of Geology
SUNY-Potsdam
Potsdam, NY 13676;
D. Jeffrey Over,
Dept. of Geological Sciences
SUNY-Geneseo
Geneseo, NY 14454;
Carlton E. Brett
Univ. of Cincinnati
Cincinnati, OH 45221
The Genesee Group succession in western New York represents deposition in dysoxic to near-anoxic settings in the subsiding Appalachian foreland basin (Fig. 1) that records a series of widespread sedimentary and biological events. We will focus on lower and medial Genesee divisions that record changes in the basin that range in age from latest Middle Devonian into the early Late Devonian. In particular, we are interested in the genesis of styliolinid (pelagic) limestone units and associated bone-conodont beds. This paper will examine the results of recent mapping and biostratigraphic work on a succession of units in the Genesee Group, ranging, in upward succession, from basal Geneseo black shale facies into the lower part of the West River Formation. We will also highlight the discovery of two mappable discontinuity-related bone-conodont beds in the lower part of the Genesee succession. The most important part of this project is the study of the regional character and inferred genesis of the North Evans Limestone (a famous conodont-bone lag unit above a major discontinuity) as well as that of the Genundewa Formation, a distinctive layer of pelagic, styliolinid carbonate that overlies the North Evans. We will examine closely the relationship of the Genundewa to both the North Evans lag deposit and overlying basal beds of the shaley West River in the context of inferred flexural and eustatic events within the basin. Another focus of this project is to better elucidate the complex chronostratigraphic story recorded in this stratigraphically condensed succession.
The change from Middle Devonian (Late Givetian), neritic shale facies of the Windom Member (Moscow Formation) of the Hamilton Group into black, condensed, basinal deposits of the lower Genesee Group [latest Givetian and earliest Late Devonian (basal Frasnian)] is abrupt and profound (Fig. 2). This change, across the widespread top-Windom, Taghanic disconformity in western New York was what initially lured two of us (Baird and Brett) away from fossil-rich Hamilton deposits into a mysterious parallel universe of black mud deposition and hypoxic water mass chemistry. The study of the genesis of exotic facies such as the diachronous, detrital pyrite deposits of the Leicester Pyrite along the disconformable base of the Geneseo black shale marking the base of the Genesee Group (see Fig. 2) was critical in understanding the genesis of black shale-roofed unconformities (Baird and Brett, 1986, 1991; Baird et al., 1989). Although work relating to the genesis of such contacts was directed to basinal units elsewhere in succeeding years, detailed study of the Genesee Group above the Taghanic unconformable contact has continued. Detailed biochronostratigraphic studies of Genesee divisional units are also ongoing. Application of new advances in sequence stratigraphy and recognition of the importance of basin flexural tectonics in the evolution of the Devonian foreland basin (Ettensohn, 1987, 1994, 1998), make the condensed western New York Genesee Group succession ideal for the testing of models (see Figs 1, 2).
Herein, we will examine several component units and boundaries within the lower and medial parts of the Genesee Group in Erie County and western Genesee County (Figs. 4-9). Some of these are newly recognized owing to renewed mapping work across this area. In particular, we will focus on regional aspects of, and inferred origin of the medial Genesee North Evans Limestone and overlying Genundewa Formation (Figs. 8-14). The first part of the paper summarizes the succession of divisional Genesee units based on new stratigraphic work across Erie and western Genesee counties. The second part focuses on new discoveries in the North Evans-Genundewa-basal West River shale succession that bear on the genesis of this distinctive, time-rich interval. This paper represents a tribute to the pioneering work of Dan Sass, who, 45- 47 years ago, traced the Genundewa Limestone across western New York and measured many of the sections described herein as part of his Masters Thesis project at the University of Rochester (see Sass, 1951). Dan, who served in World War II, was active in the fields of Environmental Science and Limnology and was instrumental in establishing both a Geology undergraduate major and an Environmental Studies program at Alfred University. He died this past March.

B3: AN EARLY LATE DEVONIAN BONE BED-PELAGIC LIMESTONE SUCCESSION: THE NORTH EVANS-GENUNDEWA LIMESTONE STORY

Cross Cazenovia Creek.

Cross south branch of Smoke Creek.

View of Lake Erie to the right.
Bluff exposure is Ledyard Member of Ludlowville Formation (Middle Devonian).

Bridge over Eighteenmile creek.
Upstream cutbank to the left shows the Middle Devonian Hamilton Group divisions unconformably overlain by Late Devonian strata of the Genesee Group.

STOP 1. Taghanic disconformity
Taghanic disconformity exposed in small cut (Loc. 2) along Lake Shore Road opposite parking area (see Fig. 9b). Exit vehicles and proceed across Lakeshore Road to roadcut slightly south of the parking lot entrance. People must watch for traffic coming down the hill around the blind comer.
This roadcut displays, in upward succession, the upper medial part of the Windom Member of the Moscow Formation (Fig. 9a), the Taghanic disconformity, a thin, local expression of the North Evans lag deposit below the Genundewa Formation (Fig. 9c), a thin unnamed and very localized dark gray to black shale unit within the Genundewa Formation, and the resistant ledge of styliolinid limestone representing the upper division of the Genundewa Formation (Fig. 9d). The disconformity below the North Evans spans possibly 7 conodont chronozones representing approximately 2 million years of geologic time (see text); units observed in the Finger Lakes region and in eastern New York (topmost parts of Moscow Formation, Tully Formation, and lower half of Genesee Group succession) are missing here such that Late Devonian strata belong to the upper part of early Frasnian MN Zone 2 are juxtaposed onto Middle Devonian beds belonging to the Middle varcus chronozone (see text).
The North Evans is here represented by a thin, 1-2 cm-thick lag unit rich in pelmatozoan debris, glauconite grains, a conspicuous mixed conodont fauna, and fish teeth, including numerous small cladodont shark teeth and rare ptychtodont tritors (Fig. 9c). Secondarily oxidized (limonitized) detrital pyrite grains are common in this mix, especially at the base. The tubular clasts represent pyritized burrows that were exhumed from the underlying Windom Shale (see text). Conodonts at the bedding-plane contact with the Windom include Ancyrodella recta that indicates the upper part of Frasnian conodont MN Zone 2. Above the North Evans is a 28 cm-thick fissile dark shale unit that can be accessed below the Genundewa Limestone overhang. As noted in the text, this shale is extremely local in extent, being found only in the vicinity of the Lake Erie shore (Fig. 9). Because there is no lag deposit and implied discontinuity separating the shale from the overlying Genundewa upper division limestone bed at this section and at Pipe Creek, we believe that the shale unit is a local basinal subfacies of the upper Genundewa division or a post-North Evans Genundewa division not normally expressed elsewhere (see text). The presence of the dark shale unit roofing the North Evans in this section is believed to explain, in part, the thin, pyrite-rich character of the North Evans lag deposit in sections where it is present (Fig. 9). More intense dysoxia associated with the basinal subenvironment represented by onset of black shale deposition could explain local dissolution of the carbonate fraction normally seen in the North Evans (see text). The overlying ledge of the upper Genundewa division (Fig. 9d) displays the typical outcrop character of this unit. This bed is composed almost exclusively of uncompressed Styliolina fissurella with accessory occurrences of the small bivalve Pterochaenia, wood debris, and rare goniatites, and it probably represents an offshore dysoxic regime characterized by prolonged sediment starvation and surface water productivity supporting Styliolina as plankton.

STOP 2. Large cutbank on west side of Cazenovia Creek (Loc. 11)
Displaying Genundewa Formation and associated units (see Fig. 12). Exit vehicles and proceed across private driveway and yard. Continue along graded trail down to floodplain of Cazenovia Creek. Turn left (north) and follow improvised trail through tall brush to outcrop approximately 300 meters north of the base of the graded trail. The improvised trail may be wet and muddy and Cazenovia Creek may be high making examination of the outcrop difficult.
This large, clean exposure shows several units to advantage (Fig.12 ); these are, in ascending order, the Amsdell Bed (Emanuella praeumbona - rich interval) of the Windom Member (Fig. 12a), the Leicester Pyrite of the Geneseo Formation (Fig. 12b), an undifferentiated interval of black and dark gray shale corresponding to the lower part of the Genesee Group (Fig.12c ), a band of laterally separated large concretions corresponding to the "grotesque concretion" level (Fig. 12d), the North Evans lag deposit (Fig. 12e), a thin zone of nodular styliolinid limestone above the North Evans comprising the lower part of the Genundewa upper division (Fig. 12f), the massive, main part of the upper division of the Genundewa Formation (Fig.12f), a thin conodont- and glauconite-bearing lag layer (Fig.12g), and the West River Formation (Fig. 12h).
This section offers an excellent opportunity to view Leicester pyrite lenses along the Taghanic disconformity if the water is not too high. These lenses (Fig.12b), laterally separated by bank intervals lacking lag pyrite, may be profiles of channel fill deposits that accumulated on the Taghanic dysoxic erosional substrate (see text). They are composed of gravel-sized grains of detrital pyrite that correspond to reworked diagenetic pyrite derived mainly from the Windom Shale. Pyritic burrow tubes make up most of the Leicester clasts, but pyritic steinkerns of orthoconic and goniatitic cephalopods, as well as fish bones and teeth, can be seen.
The black-dark gray shale succession between the Leicester and North Evans Limestone beds (Fig.12c) represents the feather edge of a much thicker Geneseo Formation succession to the east (Figs. 4-6). At the north branch of Smoke Creek (Loc. 10), four miles west of here, the sub-North Evans disconformity is juxtaposed directly onto the Windom Member.
The North Evans lag (Fig.12e) contains a mixture of pelmatozoan debris, conodonts, glauconitic grains, reworked concretions, and fish bone/tooth debris. It grades upward into styliolinid grainstone facies of the overlying Genundewa Limestone. The Genundewa is quite thick at this section; it thins dramatically eastward from here, being essentially absent at Buffalo Creek (Loc. 12), the next section to the east (see Fig. 8 A,B). The top of the Genundewa at this section is marked by a 1.0 cm-thick lag unit (Fig.12g) rich in conodonts and glauconitic grains that may link to the "Huddle Bed" diastemic horizon recognized by others (see text). This discontinuity may be, in part, responsible for erosional overstep of portions of the Genundewa Formation seen in sections east of here (see text).

Northrup Road crosses Cazenovia Creek.
This classic fossil locality can be seen to the right; the Tichenor Limestone, basal division of the Moscow Formation, holds up the low waterfall and overlies fossiliferous shale of the Wanakah Member of the Ludlowville Formation. In the distance, the high bank displays the Windom Member and overlying Genesee Group strata.

Bullis Road crosses Buffalo Creek.
Basal limestone divisions (Tichenor Member, Menteth Member) of the Moscow Formation visible below bridge in floor of creek.

Clinton Street bridge over Cayuga Creek.
Excellent exposure of Windom Member-West River succession along creek upstream from bridge. This section (Loc. 14a) has been used for field stops in the past, but access is a problem presently. We will, instead, employ a nearby section for STOP 3.

STOP 3. Upper Windom Member-into-West River Formation
Succession exposed along west-flowing tributary of Cayuga Creek (Loc. 14b; see Fig. 13). Exit vehicles and proceed west across private yard and sloped wooded area behind yard to creek bottom. Proceed left (west) down the creek to a low waterfall where the Moscow Formation-Genesee Group contact can be seen.
At the waterfall black shale deposits of the Geneseo Formation rest abruptly on the Taghanic Disconformity which caps the sloped exposure of gray Windom Shale in the falls face. Lenses of Leicester Pyrite (Fig.l3a) can be spotted owing to the trail of rust seepage developed below them on the outcrop face. The Geneseo Formation is marked mainly by a calcareous black shale interval (Fig.13b) that is capped by a discontinuity (Fig.l3c) characterized by comminuted fossil debris and a few fish bones and conodonts. It may correlate to the top-Fir Tree discontinuity in the Finger Lakes region, but this idea has yet to be tested (see text). Strata, probably equivalent to the Lodi Limestone interval in the Finger Lakes area (Fig.13d), are cryptic in this area; we will skip over this part of the section. However, above a band of prominent, brown weathering concretions (Fig.l 3e), is a black shale-on-black shale discontinuity (Fig.13f) that is marked by a pavement of flattened Styliolina in association with conodonts and cuspate cladodont shark teeth. This contact can best be accessed at the minor falls just above the concretion band where it occurs 12 cm below a 3 cm-thick styliolinid bed (Fig.13g) which is the local condensed expression of the Elevenmile Creek bed. This conodont-bone-rich horizon (unit 13f) is found to correlate eastward to the Linden Bed (see text).
At a third low falls lip the North Evans-Genundewa succession (Fig.13i,j) is observed to overlie a conspicuous band of irregular "sculptured" concretions herein referred to as the grotesque concretion layer (Fig 13h). These concretions are interpreted as underbed concretions which formed in a zone of fluctuating redox conditions below the downward migrating top-Penn Yan erosion surface; concretions of this type that formed earlier can be seen as reworked clasts within the overlying North Evans lag deposit (Fig.13i). The North Evans is well developed here and can be easily sampled. The massive, nodular, falls-capping, upper Genundewa division ledge is also well developed at this locality (Fig.13j). A thin conodont- and glauconite-rich layer, herein interpreted to be the equivalent of the Huddle Bed, is present 10 cm above the Genundewa on nearby Cayuga Creek (Loc. 14a). However, on this creek that bed is very thin and cryptic; it probably is expressed as a thin styliolinid-rich layer (Fig.13k) 12 cm above the top of the Genundewa. The West River Formation is best viewed in the north-facing cutbank section immediately upstream from the Genundewa falls. Here it consists of thin alternations of gray and black shale with a few thin siltstone beds in the upper part (Fig. B).

County Line Road crosses Durkee Creek
County Line Road crosses Durkee Creek, a tributary of Spring Creek. This shows a complete Windom Member-basal West River Formation succession (Loc. 15). Just beyond view is the thickest Leicester pyrite lens known with a maximum thickness of about 28 cm.

Darien Lakes State Park visible to north.

Bridge over Elevenmile Creek.
Long Devonian section on this creek extending from the lower part of the Ludlowville Formation up through the study interval, into the Middlesex Formation. When foliage is absent, the Tichenor and Menteth members of the Moscow Formation can be seen to cap a waterfall visible from the road.

Cross Murder Creek in Darien.
Intermittent Moscow, Geneseo, Penn Yan, and Genundewa formation sections occur along this Creek (Loc. 17).

Bridge over Little Tonawanda Creek.
The Kashong Member-Windom Member boundary is visible in the creek floor to the north of the bridge. Good Genesee Group sections are developed much further upstream near Linden. Structural dip observed in the vicinity of this creek is the local expression of the Clarenden-Linden fault zone.

Exposure of Genundewa Formation
Exposure of Genundewa Formation exposed below the Bethany Center Road Bridge on both sides of US 20. We will examine this section momentarily as part of STOP 3.

STOP 4. Genundewa Formation exposure (Loc. 23)
On both sides of US 20 below Bethany Center Road and loose Genundewa blocks in road fill along US 20 immediately east of the Old Telephone Road junction with US 20 at Bethany, New York (Fig. 14). Exit vehicles and proceed west uphill along US 20 to the Bethany Center Road overpass. The Genundewa Limestone is displayed in the protected area directly under the bridge.
This section illustrates the twofold internal division of the Genundewa Limestone, characteristic of this area, with the lower and upper Genundewa styliolinid units (Fig.14, units c and e) separated by a complex, nodular, internal zone of conodont-rich lag debris corresponding to the North Evans Limestone (Fig.14, unit d). Of particular note is the presence of Trypanites (borings) into attached Penn Yan concretions (Fig.14a); this indicates the presence of a sub-Genundewa discontinuity (Fig.14b) in this area that is distinctly older than the intra-Genundewa North Evans contact (see text). Also important is the excellent development of cephalopod limestone facies within the lower Genundewa division (Fig.14c). We will see this facies to advantage in loose blocks in the US 20 road berm downhill past the Old Telephone Road exit where the vehicles are parked. This is one of the best examples of cephalopodenkalk facies to be seen in North America. Such facies is interpreted to represent deeper water dysoxic conditions where remains of pelagic (planktonic and nektonic) organisms accumulated in a sediment starved setting (see text). The goniatite fauna includes Koenenites styliophilus, Acanthoclymenia genundewa, and Tornoceras uniangulare compressum (see Kirchgasser et al., 1994, for illustrations of Genesee Group goniatites). Pyritic replacements of Acanthoclymenia occur here at the contact with the overlying West River Formation. Proceed east, downhill along US 20 to road fill berm past the Old Telephone Road pull off. Sample blocks exposed on the partly overgrown berm slope.

B4: THE PENN DIXIE SITE: A CLASSIC AND UNIQUE PALEONTOLOGICAL & OUTDOOR EDUCATION CENTER
Jerold C. Bastedo
Executive Director
Penn Dixie Paleontological and Outdoor Education Center
P.O. Box 772
Hamburg, New York 14075
www.penndixie.org
The Hamburg Natural History Society, Inc. (HNHS) is a nonprofit educational corporation that owns and operates The Penn Dixie Paleontological and Outdoor Education Center in Hamburg, New York (Fig. 1). The HNHS was founded in 1993 to promote the study of the natural sciences, with a particular emphasis on field activities associated with the geological and biological sciences. The HNHS offers a wide variety of hands-on educational programming to students of all ages, both at the Penn Dixie Site and off-site at local schools, libraries, and civic group meetings. Since its inception, the HNHS has expanded its educational curriculum to include public educational programming in astronomy and ornithology to complement its core study in geology and fossil collecting and identification. Unlike conventional museums or research facilities, the Penn Dixie Site is a hands-on outdoor educational facility - one at which visitors of all ages are encouraged to actually collect and keep 380-million-year-old fossils - "Where Science Comes Alive".
The site of a former quarry operation that was the source of calcareous shale excavated and used for cement aggregate by the Penn Dixie Cement Company. A majority of the 57-acre site was quarried until the late 1960s, during which time 9 to 10 feet of shale was removed from the surface. A gray, somewhat flat "desert-like" or "lunar landscape-appearing" surface now occupies a majority of the site. After quarry operations ceased, weathering forces began to expose 380 million-year-old Devonian fossils preserved within the Windom Shale. This highly fossiliferous unit underlies the entire site and provides an inexhaustible supply of fossils. In addition to the Windom Shale, several limestone units (the Genundewa, North Evans, and Tichenor) outcrop on the surface. The Wanakah Shale is also exposed, underlying the Tichenor Limestone, in a tributary that flows into Rush Creek and in cliffs along Rush Creek on the northern section of the site. All of these units contain a variety of fossils.

B4: THE PENN DIXIE SITE: A CLASSIC AND UNIQUE PALEONTOLOGICAL & OUTDOOR EDUCATION CENTER

STOP 1. Penn Dixie Paleontological and Outdoor Education Center

B5: USING MARINE FOSSILS TO UNLOCK THE MIDDLE DEVONIAN PALEOENVIRONMENTS OF WESTERN NEW YORK (FOR K-12 TEACHERS AND COLLECTORS)
PHILIP J. STOKES
SUNY at Buffalo
pjstokes@geology.buffalo.edu
JAMES J. ZAMBITO, IV
University of Cincinnati
Cincinnati, OH 45221
zambitjj@email.uc.edu
The physiographic province of the Allegheny Plateau offers a geologically rich visage of the Paleozoic prehistory of western New York State. The region possesses a relatively complete (and highly fossiliferous) section of Middle Devonian rocks. This section of Devonian strata thickens towards the east, directly reflecting the resultant sedimentation from the Acadian orogeny, and inspiring the name of the Catskill clastic wedge to the strata (Brett, 1986). As mountain building progressed to the east and southeast, most of New York State was covered by a shallow sea in what is known as the Appalachian Basin (Isachsen et al., 2000). During this time, the amount of sediments being deposited changed, the type of sediments changed, and local sea levels rose and fell. These paleoenvironmental fluctuations set the stage for major changes in the regional paleontology during the Middle Devonian.
This paper and the accompanying field trip seek to introduce science educators and fossil collectors to some interesting Devonian sites in western New York (Figure 1). The trip explores rocks chiefly from the Middle Devonian Hamilton Group, which represents a 5 to 7 million year time period between about 377 and 384 million years ago (Brett, 1986). The Hamilton Group includes four formations: the Marcellus, Skaneateles, Ludlowville, and Moscow. Formations within the Hamilton group are divided by three separate limestone units: the Stafford, Centerfield, and Menteth/Portland Point (Patchen and Dugolinsky, 1979; Brett, 1986). The Genesee Formation, which in some places overlies the Hamilton Group, is another important unit in understanding the regional geologic setting (Brett and Baird, 1994). The sequence of Devonian depositional events along with data from fossil assemblages helps us to characterize the "big picture" of the local geological history.

B5: USING MARINE FOSSILS TO UNLOCK THE MIDDLE DEVONIAN PALEOENVIRONMENTS OF WESTERN NEW YORK (FOR K-12 TEACHERS AND COLLECTORS)

Parking for Stop 1.

STOP 1. Retsof Locality (of older literature).
Cross Rt. 63 at the entrance to the parking lot. Proceed down the gravel trail (westward) towards and just beyond the Rt. 63 Retsof Switch Structure. Head south about 250-300 meters until the flat bedding surface turns into outcrop near old railroad yard. If vegetation is high, it may be difficult to find the trail. The best collecting from this site is found along the bluff wall adjacent to the rail yard, although surface collecting is also possible in some places.
Geologic unit: Kashong Shale Member of the Moscow Formation, Hamilton Group
Geologic age: Middle Devonian
Common fossils: Crinoids, brachiopods, corals, bryozoans

STOP 2. Alden Pyrite Beds
(east-facing shale pit on Bethany Center Road).
Exit vehicle and walk approximately 150 meters west into the clayey hills to the right (west) of the road. Pyrite beds outcrop from the weathered bedrock surfaces near recently incised channels throughout the site. Calcareous concretions eroding from the outcrop often contain brachiopods and cephalopods. Pyritized fossils also appear on muddy flats below weathered outcrop.
Geologic unit: Ledyard Shale Member of the Ludlowville Formation, Hamilton Group
Geologic age: Middle Devonian
Common fossils: Pyritized brachiopods, trilobites, rugose corals, cephalopods, gastropods

STOP 3. Delaware, Lackawanna, and Western Railroad Tracks.
From Stop 2, head back to the road and walk north about 200 meters, past the guard rails on Bethany Center Rd. and descend left (west) on gravel path into the railroad bed. Abundant fossils can be found both on the surface and in weathered bedrock fragments. Possible lunch stop.
Geologic unit: Centerfield Limestone Member of the Ludlowville Formation, Hamilton Group
Geologic age: Middle Devonian
Common fossils: Rugose and tabulate corals, bryozoans, brachiopods, trilobites, crinoids

STOP 4. Cayuga Creek.
Cross Eastwood Rd. and descend to the creek via the dirt trail alongside the overpass. Take note of the bedrock on which the overpass trusses are built. The black shales of the lower Geneseo are nearly void of fossils at this location except for sporadic pavements of chonetid brachiopods. Lenses of the Leicester Pyrite contain pyritized burrows and fossil steinkerns.
Geologic units: Windom Shale Member, Moscow Formation, Hamilton Group; Leicester Pyrite (in lenses along unconformity), Genesee Formation; and Geneseo Shale Member, Genesee Formation.
Geologic age: Middle and Upper Devonian
Common fossils: No common fossils, sparse brachiopods

STOP 5. Buffalo Creek.
After parking, descend down ramp trail and cross old Buffalo Creek bridge. Proceed down one of several trails to the creek. Approximately 300-400 meters downstream from the bridge, cross the creek in a shallow place to get access to the vast exposures of bedrock. The Lower Windom Shale is particularly accessible on the north bank of the creek upstream of the old Bullis Rd. bridge.
Geologic units: Wanakah Shale and Jaycox Shale Members, Ludlowville Formation, Hamilton Group; Tichenor Limestone, Deep Run Shale Member, Meneth Limestone, Kashong Shale Member, and Windom Shale Member, Moscow Formation, Hamilton Group; and Geneseo Shale, Genesee Formation.
Geologic age: Middle and Upper Devonian
Common fossils: Trilobites, brachiopods, corals, bryozoans

B6: QUATERNARY GEOLOGY AND LANDFORMS BETWEEN BUFFALO AND THE GENESEE VALLEY
R.A.Young
Department of Geological Sciences
SUNY College at Geneseo
Geneseo, New York 14454
and
Jason P. Briner
Buffalo, New York 14260
The Quaternary geology and landforms in western New York State north of the southern Finger Lakes (Valley Heads moraine) formed during the final recession of the Late Wisconsin ice sheet from its terminal position (Figure 1). There are a few places where Middle Wisconsin drift is preserved, such as in the Genesee Valley (Young and Burr, 2006), that indicate there is a more complex, older glacial record yet to be explored. This field trip explores the dominant geomorphic features created by the final stages of ice recession across a scarp-dominated bedrock landscape from Batavia to the southern Lake Ontario shoreline.
The complex sequence of proglacial lakes that accompanied glacial recession across the Great Lakes region is well documented in the geologic literature of the last century, but some events continue to be subjects of detailed study and debate. For the purposes of this field trip, the creation of glacial Lake Whittlesey (Figure 2) at the time of the Port Huron advance (circa 13,000 14C yrs BP) is an appropriate place to begin a discussion of the major events in western New York that culminated in the formation of early Lake Ontario, Lake Tonawanda, and the development of Niagara Falls in the Niagara Gorge.

B6: QUATERNARY GEOLOGY AND LANDFORMS BETWEEN BUFFALO AND THE GENESEE VALLEY

STOP 1. Buffalo/Niagara Falls moraine complex
The Buffalo/Niagara Falls moraine complex. Gravel quarries are cut into a large outwash fan (labeled Wisconsin outwash gravel (Wog) on Muller (1977) which borders the Buffalo Moraine. Borings made for the study of the gravel aquifer in this area show two till units bracketing up to 40 ft. of sand and gravel, indicating that the outwash fan pre-dates a more recent ice advance. Thin red lake clay partings separate thick rippled sand units, indicating that the outwash sand/gravel was likely deposited subaqueously. The overlying till and sand/gravel unit shows evidence of flow and deformation. The till has been traced at least to the Alden Moraine (Figure 3), 8 km to the south.

Cross Crittenden Rd. and onto Lake Warren beach ridges
Glacial Lakes of the Erie Basin. Because the Laurentide Ice Sheet (LIS) was sourced in the north (Figure 1 ), and Lake Erie drains north, large pro-glacial lakes were created as the LIS flowed into and out of the Lake Erie basin. The highest lake level recorded in the greater Buffalo region, termed Glacial Lake Whittlesey (Figure 2), created mainly an erosional escarpment between 850 and 900 ft. asl (260-274 m asl). Present Lake Erie elevation is 570 ft asl (174 m asl). As the LIS retreated to the Alden, Buffalo, Niagara Falls and Batavia moraines (Figure 3), the lake level dropped to a new level termed Glacial Lake Warren, which deposited beaches along the field trip route from 820 to 850ft. asl (250-260 m asl) between 12,700 and 12,500 14C yr BP. Following Glacial Lake Warren, lower lakes termed Grassmere, Lundy, Algonquin and Dana existed; these lakes pre-date non-glacial Lake Tonawanda (see below), which existed from slightly before 12,000 14C yr BP up to ∼9000 14C yr BP. See shoreline elevations and relative amounts of postglacial uplift on Figure 8.
Key chronological constraints on the glacial history of western New York (from Muller, 1977; Calkin, 1970)_.
14,100 to 13,400 | Valley Heads, Lake Escarpment, Gowanda Moraine deposits
13,000 | Lake Whittlesey, Hamburg Moraine
12,700 | Lake Warren (high stand), Alden Moraine
12,500 | Lake Warren (low stand), Batavia Moraine, Niagara Falls Moraine
12,400 to 9,000 | Existence of Lake Tonawanda
12,200 to 11,000 | Lake Iroquois (main phase)

This road closely follows the Batavia moraine.
The Batavia moraine near Batavia has noticeably greater and more irregular relief than the moraines that will be visited to the north and west. This is due, in part, to the formation of some of the younger moraines at a submerged ice front (Figure 7). In this environment moraines are subject to being covered by lacustrine sediments and then to being winnowed and eroded by waves as lake levels fall. The west end of the Batavia moraine (930+ ft as1) projects above the elevations of Lake Warren, and thus it would have escaped wave erosion when an eastern finger of Lake Warren formed a narrow extension into the Genesee Valley with shorelines at current elevations close to 860 ft asl near Geneseo, NY. While driving along this Batavia moraine segment, note its dimensions and relief for comparison with the more subdued Barre moraine to be visited later (Figures 3, 5). According to Calkin (1970) the Batavia moraine is believed to have formed around the time of the lowest Lake Warren stage, which coincided with ice recession from that position (Figure 3). Lake Warren had three phases, Warren I and II are higher and older, Warren III is the lowest and youngest, and followed a short-lived drop to a "Lake Wayne" stage (Figure 8).

Route crosses crest of Batavia moraine.

STOP 2. Onondaga Limestone Outcrop
Traveling North, Maple Rd. drops off the Onondoga Escaprment. Old quarry and recycling center exposes Onondoga limestone (Middle Devonian). Note how, in general, the regionÕs resistant bedrock units lead to relief (30-100 ft. escarpments, Figure 12 ) that dwarfs the glacial topography. STOP 2. Examine the nature of the resistant Onondaga Formation.

Crossing Nichols Hill drumlin.
Note height and trend of this prominent drumlin.

Begin ascent of drumlin along its axis.

STOP 3. Tonawanda Lowland.
Pullout on right side of road affords nice view of drumlin relief and inter-drumlin bog (Figure 13). The two drumlins here, trending WSW, show the direction of ice flow during the late stages of the last glaciation.
This location is approximately one mile north of the southern edge of Lake Tonawanda (note surrounding wetlands), which was nearly 7 miles wide at this longitude. Lake TonawandaÕs northern limit was determined by the rising dip slope of the Lockport Escarpment. The southern limit was formed, in part, by moraines (Figure 3) and by the gradual increase in elevation toward the Onondaga Escarpment (Figure 7). Lake Tonawanda formed following the lowering of Lake Warren toward the Lake Iroquois level, which is north of the Lockport Escarpment (Figure 3). Lake Tonawanda had several northerly outlets (Figure 3, Arrows), which are progressively older and higher toward the east. Lake Tonawanda waters flowed off the escarpment into Lake Iroquois, when it first formed. As shown on Figure 2, the easternmost outlet formed first and was succeeded in turn by the more westerly outlets as indicated by letters A through E on Figure 3. Current threshold elevations near the eastern (Holley outlet) are near 650 feet, whereas the threshold elevations near Gasport are closer to 600 feet. The westerly displacement of the outlets was dictated by the greater amount of postglacial rebound toward the northeast (Figure 8). Eventually the Lake drained completely after the outflow was shifted toward the modern position of the Niagara River.

Veer left as road descends northern drumlin flank.

Note stony till just south of intersection.
Road crosses Barre moraine.

STOP 4. Eagle Harbor Sand and Gravel quarry. (Lunch)
Part of the Burma Woods esker and kame moraine complex. Dozens of small-scale quarries and this large one in this vicinity.
Burma Woods Esker. This esker and outwash complex (seen in a modern excavation) may have formed during the earliest stage of Lake Tonawanda, when the ice had just retreated from the Barre moraine (Muller and others, 2003). Muller and others (2003) note that the position of this and similar eskers indicates early southward meltwater flow into Lake Tonawanda. Although the esker deposits are not currently visible in the modern pit, it is possible that the deformed sediments being excavated are contemporaneous with an ice position fronting on early Lake Tonawanda. The sediments visible in August 2006 at the working face (Figure 15) indicate that saturated sediments were deformed by intraformational slumping, and possibly by upward groundwater discharge immediately in front of the ice margin. This will be the longest stop on the trip (including lunch) and will allow time for discussion of how the currently exposed sedimentary structures may relate to the history of ice recession from the Barre moraine, or to the development of Lake Tonawanda.

STOP 5. Burma Woods esker
Turn left onto Maple St. and pull off to the right at one tenth of a mile to see the roadcut through the Burma Woods esker. This esker is only several meters high, but contains well-rounded sand and cobbles with some boulders. Watch for traffic.

This road is built atop the Barre Moraine.
Note how the Burma Woods esker trends perpendicularly to the Barre Moraine (Figures 3, 5).

STOP 6. Lake Tonawanda spillway.
This is the location of one of the main spillways from Lake Tonawanda across the Lockport Escarpment south of Medina (see Figure 3 and discussion for Stop 3). At this location the lake outflow was able to cut down through the Barre Moraine, probably beginning at a low saddle in the moraine crest.

Cross Lake Tonawanda outlet channel.

STOP 7. Victor Fitchlee Park
Rest stop and short walk into Royalton Ravine (lake outlet at East branch of Eighteenmile Ck). This Park has restroom facilities and is located adjacent to a deep outlet channel passing through Gasport. A short walk down Royaltine Ravine leads to the broader Tonawanda channel, and the Ravine exposes shaley bedrock, which indicates how thin the glacial drift is at this location.

View to left off of Niagara Escarpment.
Smoke stack in distance on shore of Lake Ontario, 12 miles away.

Road follows old Earie Canal

Note series of locks to get the Erie Canal across the Niagara Escarpment.

STOP 8. Outwater Park.
Overlook to Lake Ontario from top of Niagara Escarpment.

STOP 9.Fossiliferous surface of the Lockport Dolomite.
Hidden driveway on right leads to footpath to old quarry site where glacial drift has been stripped away revealing a striated fossiliferous surface of the Lockport Dolomite.

Good exposure on left of Medina Group
Good exposure on left of Medina Group in Niagara Escarpment. Alternate STOP 9.

Glacial lake Iroquois.
Route 104 follows the glacial Lake Iroquois beach.
Glacial Lake Iroquois is probably the best known of the proglacial lake stages in the Ontario basin, because of its prominent relief and the location of the original Route 104, which followed its course along much of the lake shore. The development of the Niagara Gorge and the retreat of Niagara Falls, and their relationship to Lake Tonawanda, are complex and interesting stories, described recently by Brett and Calkin (1996) in the guidebook for the Northeast Section Meeting of the Geological Society of America in Buffalo.
Lake Iroquois formed around 12,200 14C yrs BP, and persisted for several hundred years, based on existing chronology and the strong development of its shoreline. Many of its features are as well developed as those seen along the modern shoreline of Lake Ontario. It is not generally realized that Lake Ontario subsequently reached a much lower level when the ice sheet finally vacated the north flank of the Adirondacks (Covey Hill outlet) and Lake Iroquois waters escaped to the Atlantic Ocean (glacially lowered) via the St. Lawrence Valley. Since reaching this lower level, which was between 200 and 300 feet lower at Rochester, NY, the modern lake has been steadily rising to its present level. This is the result of postglacial tilting of the basin and uplift of the threshold that controls flow out through the St. Lawrence River. It was previously thought that Lake Ontario was a marine or brackish arm of the ocean for a short period (Champlain Sea), but modern studies have demonstrated that marine waters did not advance much west of Ogdensburg, NY, during that post-Iroquois, early Lake Ontario low stand. Further reading about the glacial history of the Ontario and Erie basins is contained in Geological Association of Canada Special Paper 30, Quaternary Evolution of the Great Lakes (Karrow and Calkins, 1985), and in Special Paper 35, The Late Quaternary Development of the Champlain Sea Basin (Gadd, 1988).

Road is still on a beach berm
Here, the road is still on a beach berm, but this beach is a spit-like feature that diverges westward from the main Iroquois beach strand.

STOP 10.
Stop to see sand and gravel of the Iroquois beach.

You immediately drive off Iroquois beach
You immediately drive off Iroquois beach and onto flat Iroquois lake plain.

STOP 11. Wilson Tuscarora State Park (time permitting).
Compare character of modem Ontario beach and its sediments with those most recently viewed at Stop 10.